segment boundaries, the 1894 ruptures and strain …

Pergamon

.I. Geodynnmics Vol. 26, No. 24, 461486, 1998 pp. 0 1998 Published by Elsevier Science Ltd

All rights reserved. Printed in Great Britain PII: SO264-3707(!W)ooo66-5 0X&3707/98 $19.00+0.00

SEGMENT BOUNDARIES, THE 1894 RUPTURES AND STRAIN PATTERNS ALONG THE ATALANTI FAULT, CENTRAL GREECE

ATHANASSIOS GANAS,‘* GERALD P. ROBERTS2 and TZETTA MEMOU3

‘Department of Geography, University of Reading, Whiteknights, Reading RG6 6AB, U.K. *The Research School of Geological and Geophysical Sciences, Birkbeck and University College London, Gower Street, London WClE 6BT, U.K. ‘Division of Geophysics, IGME, 70 Messoghion Str., 115 27 Athens, Greece

(Received 27 April 1997; revised 2 June 1997; accepted 5 June 1997)

Abstract-The Atalanti Fault is a large active normal fault segment inside the Gulf of Evia Rift system (Central Greece), that last ruptured during the April 1894 earthquake sequence. Using structural and geomorphological interpretations of digitally processed Landsat TM satellite imagery, two regions of i) low topography, ii) minimum hinterland development and iii) transverse bedrock ridge development, 34 kilometres apart were identified; these regions are suggested to be segment boundaries constraining the length of the fault. From throw profiles and displaced syn-rift strata, we estimate a minimum slip of 810m at the central region of the fault (Tragana), increasing to a value of 1200 meters within the Asprorema embayment area. These figures averaged over a time span of 3 million years (age of oldest offset syn-rift), yield mean slip rates of at least 0.27 to 0.4 mm/year.

Field studies were also conducted along the length of the Atalanti Fault Segment to re- examine and map the 1894 ruptures. The surface break is only preserved locally where the footwall comprises a resistant bedrock lithology (limestone), whilst the rest of the rupture noted in historical records propagated along the contact with the volcanic pre-rift, as well as within the syn-rift, and has since been eradicated due to man-made changes in surface morphology. The surface breaks appear not to have crossed over the segment boundaries that we propose, but seem to have ruptured the full length of the Atalanti Fault Segment, that is, 34 km. These observations suggest that the 1894 rupture is the longest mapped within Central Greece. However, it remains unclear whether the ruptures were produced solely by the 271411894 earthquake, or by two events, one week apart. We discuss the implications for fault-behavioural models and seismic hazards for the Atalanti area. 0 1998 Published by Elsevier Science Ltd. All rights reserved

*Author to whom all correspondence should be addressed: E-mail: [email protected].

461

I. INTRODUCTION

In Central Greece (Fig. I 1. cruhtal cxtcnhion 14 ;lcc<mmvdated by slip along large normal

fault segments of planar geometry. These faults ~ODIIC! it series of E-W to NW-SE half- grabens exhibiting major. but systematic variations in topography and bathymetry along

their axes (e.g. Roberts and Jackson. 199 I ; Roberts and Gawthorpe, 1995). Syn-rift deposits

of Neogene and Quaternary age have been mapped within the continental basins of Eastern Central Greece (Evia. Phiotis. Viotia: IGME. 1965. 197X; Mettos el al.. 1992). The basins are arranged in a ‘domino-style‘ of fault-controlled continental extension (e.g. Westaway. 1991 ). More than 60”/0 of the extensional deformation is accommodated by large earth-

quakes ( > MS 5.8; Ambraseys and Jackson. 199~): Billiris r/ ul.. 1991). Published earthquake catalogues (e.g. Papazachos and Papazachou. 19X9: NEIC-PDEs) show that most historical

and instrumental seismicity is localised within the major marine basins on either side of the Gulf of Evia. These are the Gulf of Corinth (Roberts and Koukouvelas. 1996; Armijo CI

N/.. 1996) lying 60 km to the south, and the Gulf of Pagasai region in Southern Thessaly (Pavlides. 1993) lying SO km to the north. The Gulf of Evia rift has been a seismically-quiet

area. with the last surface rupturing events on 77 /4; I X94 along the Lokris coast (Skouphos.

1 X94).

Fig. 1. C;ener,d gwlogd map 01’ Easlern c cn~ral ~~reecc. ~ho~~~n~ geological formations (adapted from IGME.

1989) and reglonal neotectomc deformation (fault segments and marine basins). Pre-rift rocks are mainly Mesozoic

carbonates and ophiolites. Syn-rift is mostly composed of Pliocene lacustrme formations and Quaternary alluviA

deposits. Fault segments shown have been defined In Ganas (‘I N/. (1996). Symbol RB is Renginion basm. Box at lower left show5 a map of Greece with the region of study.

Segment boundaries, the 1894 ruptures along the Atalanti Fault, Central Greece 463

A remarkable feature of the 1894 earthquakes is that the ruptures are purported to have been 60 km in length (e.g. Richter, 1958 pages 617618; Karnik, 1971 page 104; Lemeille et al., 1977; Rondoyianni-Tsiambaou, 1984; Stiros and Rondoyianni, 1985, 1988; Papazachos and Papazachou, 1989 page 100; IGME, 1989). In contrast, earthquake ruptures mapped along fault segments in central Greece, such as those on the southern shore of the Gulf of Corinth from the 1981 and 1995 earthquakes, seem to grow up to a length of < 15-20 km (Jackson et al., 1982; Ambraseys and Jackson, 1990; Roberts and Koukouvelas, 1996; Hubert et al., 1996). Also, further north, the 13/5/95 Grevena earthquake (MS 6.6) produced surface ruptures along 8-12 km (Pavlides et al., 1995; Meyer et al., 1996) of a much larger normal fault (27 km; Clarke et al., 1997). In this paper we re-examine the extent of the 1894 surface breaks, and the length of the host fault segment using field and satellite observations. Our conclusions show that although the entire length of the fault segment we define was ruptured, a length of 34 km, it is unclear whether the ruptures were produced solely by the 271411894 earthquake, or by two events, one week apart.

Sections 24 of this work present new information concerning decametre-and-km scale structures extracted from digital processing of low-sun angle satellite images (Landsat Thematic Mapper; see Salomonson and Stuart, 1989). These structures include hinterland geomorphology, transverse bedrock ridges and cross faults, all of which can be used to infer the positions of persistent segment boundaries along the strike of seismogenic normal faults (see Crone and Haller, 1991 for methodology). In addition, data (field and stratigraphic) on the throw distribution across the Atalanti Fault are presented together with the results of electrical surveying across the Atalanti plain conducted by TzM by use of the resistivity method (see also Memou, 1986). Structural data concerning fault displacement, fault-slip directions and bedrock scarp morphology are presented to support observations of fault-segment length. Section 5 includes our re-examination report of the rupture path of the 27/4/1894 M7 earthquake (see also Skouphos, 1894, pages 442448).

2. THE GEOMETRY AND KINEMATICS OF THE ATALANTI FAULT SEGMENT

2.1. Remote sensing and geomorphic analysis Remote sensing, a relief-sensitive mapping technique can be used to infer fault segment

positions in rift systems because in such areas the configuration of relief may be linked with systematic variations of fault displacement (Anders and Schlische, 1994). For example, in the Gulf of Corinth rift system (Roberts and Koukouvelas, 1996; Armijo et al., 1996) as well as within the Basin and Range extensional province (Wallace, 1978; Crone and Haller, 1991) the surface topography is controlled by long-term growth of normal faults. The footwalls of such faults are marked by high mountains, where pre-rift rocks are exposed; the elevations of these mountains change along fault strike. Areas of relatively low elevation lie at the ends of the fault segments. These areas of low topography can be either recognised from space by the construction of high resolution digital elevation models (e.g. Ganas et al., 1996), or from a simple examination of I:100 000 or other suitable topographic maps of rift systems (e.g. Agar and Klitgord, 1995). These low-relief positions are characterised by a large cumulative deficit of seismic slip along the main rift-bounding fault, together with an increased ‘distributed deformation by smaller faults (Zhang et al., 1991), and, therefore, are primary candidates for recognition as persistent segment boundaries (PSBs).

To map the length of the Atalanti Fault Segment and segment boundaries from space,

we used a winter Thematic Mapper (TM) scene (Figs 2 and 4; scene ID = 4202208284, Path/Row: 183/33, 28.5 x 28.5 m pixel size), provided by IGME (Athens), The image was

processed by AG using a Sun Sparc20rM workstation, using the Erdas ImagineTM image

processing software. This scene is particularly suited for structural interpretations because of the low sun-angle (25 degrees) during its acquisition date (28 January, 1988) and its

geometric rectification into the UTM projection, that conforms to the Greek national cartographic grid. The low sun-angle combined with a southeastern sun azimuth in the

image selected, provide an almost ideal imaging geometry for E-W ( + 20 ) trending normal faults downthrowing to the North, such as the N290-striking Atalanti fault segment (Fig. 2). This is because relief created by the dome-shaped footwall uplift we describe below (see Fig. 5a,b; compare with Fig. 4) casts long shadows to the north of the fault segment.

Shadow widths diminish towards segment boundaries because of the low values of relief.

In addition, UTM-geocoded imagery can be used to measure geometrical properties of structures and landforms within acceptable accuracies (_+ 1OOm) for regional-scale

mapping. To provide an insight into the digital manipulation of imagery, it was found that footwall

topography along the Atalanti fault is better imaged using principal component trans- formations and linear stretching of the near-infrared spectral bands of the sensor (TM4. 0.7660.90pm: TM5. 1.55 I .75 jtm), than using other techniques such as the composition of

Fig. 2. Early-morning, 28.5 m resolution image of the Malesina Penmsula and surrounding areas. Lokris, Central

Greece. The image has been processed (Principal Component Analysis, 1st Component shown) to enhance

topographic detail. Exceptionally bright areas represent sun-facing relief almost free of vegetation cover. Black

areas are sea and shadows. The prominent linear feature running diagonally across the image is the Atalanti Fault.

AF indicates the April 1894 rupture localities where the co-seismic displacement was measured in this study. Other

lettering is explained in the text. The vertical sides of the image are oriented 9 degrees east of North.


Atalantl Faull Trace \ from TM Image f Stream .‘%tchmant Boundary 0,

4km J

Fig. 3. A line interpretation of Fig. 2 showing footwall catchments and streams of the central part of the Atalanti Fault Segment. Selected streams are also shown in the hangingwall, where they flow away from the fault. It is suggested that because of the observed shape symmetry of the fault hinterland, the fault has grown as one earthquake segment despite the existence of three bends west of the urban areas of Proskynas, Tragana and

Kyparissi, respectively.

false colour triplets and ratio-images (e.g., Sabins, 1987 pages 235-277; Bushara, 1996). In particular, structural information is most easily and clearly derived from interpreting the first principal component (e.g. Figs 2 and 4), because this technique produces an image containing the maximum variance (_ 80%) of the TM multispectral dataset (6 bands), that is determined by illumination differences resulting from variations in surface relief.

In Fig. 2, the prominent lineament running diagonally across the image is a large portion of the Atalanti Fault (see Fig. 1 for location). The lineament is of a morphologic origin (a directional break in surface slope; see HAGS, 1975; also Fig. 6) and of low sinuosity (actual fault length/straight fault line = 1.12), as expected for active neotectonic faults where slip rates keep pace with erosion rates. Notice that structural and geomorphic features predicted by segmentation theory (e.g., Crone and Haller, 1991; Leeder and Jackson, 1993; Jackson and Leeder, 1994) to occur at the ends of active faults are clear on this image (Fig. 2). These include 1) the short, bright and dark lineaments representing smaller normal faults (P for Pavlos fault; M for the Malesina fault; V for the Vivos fault); 2) the transverse bedrock ridge (TBR) to the west of Larymna Bay (a second one occurs immediately to the east); 3) the contrasting flow directions of the streams in the area (see arrows close to D). Also, footwall drainage occurs in the form of ‘wine-glass’ valleys further west; such features are known to characterise the central portions of fault segments (Wallace, 1978). Drainage

466

t-lg. 1. Landsat- image 01’1he northwestern part 01’ the Atalantl Fault Segment. Illumination ic Cram the bottom

I-lght and thin black stripe\ reprornt rectilinear \alle)\ 111 the t’~u>t\\;~ll ol‘thc l’ault The tou n of Atalanti is ahohn

hy the arrow At. Ky is an ahhrevlatton for the \illagc OF K!p:w\u and Aj f’or Asprorema. respectively. This image

also contains information on I~tholopy (Letter> I) and \’ ind~catc dolomite and volcanic bedrock. rcspectivelq).

associated with the ‘wine-glass‘ \‘alleys ih dircctcd north-south (see a line rnterpretation 01‘ this image in Fig. 3). whereas hangingwall drainage in the region lo the north and to the east of the Larymna TBR is flowing to the east. Drainage between the Malesina and Vivos faults (about four to six kilometres to the north of the Atalanti fault; Fig. 3) flows to the north towards the Gulf of Evia.

It is suggested that the distribution and relative growth of footwall catchments along isolated fault segments (the hinterland extent) can be used as a geomorphic criterion toward PSB recognition. As normal fdults grow laterally by accumulation of seismic slip (e.g. Cowie and Scholz. 1992; Jackson and Leeder. 1994). mean slip rates are lower at the ends of the faults than along their central portion. Thus. assuming identical footwall lithologies. mean erosion rates must follow the same pattern producing smaller catchment areas in the footwall in segments boundaries compared to the centres of faults. It is suggested that such organised patterns are identifiable from space. For example. compare the two areas marked AF in Fig. 2 (see also Fig. 3). Although pre-rift lithology is different (see key in Fig. 1). it is clear that fault hinterland tapers out towards the east. Similarly, the same methodology can be applied in the northwestern end of the fault (Fig. 4). The organised pattern of footwall catchments comprising a series of rectilinear ‘wine-glass’ valleys vanishes at about the geographic latitude of the town of Atalanti (symbol At). Ganas and White (1996) have also mapped 6 km to the north-west of the town of Atalanti, a transverse bedrock ridge and a region of low topography separating the Atalanti and Renginio basins (Fig. I, symbol RB), which support the location of a segment-boundary in that area.


Fig. 5. Field photographs along the Atalanti Fault. Top) View of ‘wine-glass’ valleys (see vertical arrow) in 1 .he

footwall area south of the town of Atalanti (April 1995). The view is towards the SW. The horizontal arrow points

to a mountain peak with the highest footwall elevation, 1079 m. The fault runs along the mountain front, however

its plane can be only inferred by the prominent break in slope. Bottom) View of the exhumed fault plane (October

1996) in the Asprorema area, inside the channel, 200m east (downstream) of the bridge in the road Atalanti-

Kolaka. Thick arrow points to the pre-exhumation vegetation line. Thin arrow points to a location where a fresh

stripe can be seen above the exhumation level. The fault plane seen is about 60 metres high and is striated down-

dip from repeated co-seismic displacements. The lineation data set is presented in Fig. 10.

In summary, there are three pieces of information provided by the structural and geo-

morphological interpretation of Figs 2 and 4: (a) The systematic decay in fault hinterland

development towards both the south-east and the north-west; (b) drainage flow directions (Fig. 3); (c)the recognition oftransverse bedrock ridges. It is proposed that this arrangement

of structural and geomorphic features defines the length and position of the Atalanti Fault Segment. The segment is about 34 km long, and runs between the towns of Atalanti and

Larymna. Below, we present other datasets to support our interpretation.

3. STRATIGRAPHIC DATA AND MEAN SLlP RATE

In order to test our hypothesis concerning the length and position of the Atalanti Fault Segment we studied the long-term evolution of the fault by looking at its throw distribution

profile (Figs 6 and 7). Figure 6 contains two datasets which are jointly presented in Table I. The first data set comprises the results of a geoelectrical survey across the Atalanti plain

(Memou, 1986; see Fig. 8 for location of the profiles) which defines the thickness of the

syn-rift deposits (Neogene to Recent). The second dataset (south of Kyparissi) comprises a series of elevation data for syn-rift and pre-rift formations on either side of the fault (Table 1). The pairs represent the highest footwall and hangingwall positions of a particular

stratigraphic unit and are located at selected sites across the fault where accurate elevation

Gulf of Evia

Fig. 6. Map showing the trace of the Atalanti fault (thick. light gray pattern line), together with selected elevation

contours and pairs of localities (black dots) across the fault. used to extract throw estimates. Prefix F denotes a

footwall locality, prefix H a hangingwall one, respectively. Black boxes denote occurrences of syn-rift rocks of

Pliocene age, matched across the fault plane. Black circles indicate localities where data on fault-slip directions

were collected. Thick black line labeled I (map centre), is the distance d in the diagram shown in the inset box

(upper left: see text for discussion of the diagram). Grey-filled circles in top-left and bottom-right are the segment boundary proposed positions. Cross-hatched circles are urban areas referred to in the text.


MINIMUM THROW ESTIMATES ALONG THE ATALANTI FAULT SEGMENT

SE-NW

-400

-600

DISTANCE ALONG FAULT (km)

Fig. 7. Graph showing the distribution of the Atalanti Fault throw. Thin dashed lines indicate throw profile for

syn-rift localities, thin continuous black lines for pre-rift localities. Pre-rift localities are marked by triangles and

syn-rift by squares. For each pair of points its footwall and hangingwall elevation is given by the number to its

right. Both pre-rift and syn-rift throw distributions vary systematically along the length of the fault segment and

attain their highest value between kilometres 24 and 28 (Asprorema area).

information is available (+ 1 m; triangulation bench marks from the Hellenic Army Geo- graphical Service 1:.50 000 maps). We use this information to compare the elevations of stratigraphic formations across the fault in order to assess the fault throw.

The critical assumption here is that these data points (Pl to PlO) were all lying at about the same horizontal datum surface in the pre-faulting state. If this assumption is accepted then, the throws (vertical displacements) calculated from their height differences in Table 1 are least estimates of throw at that particular area along the fault strike. True throws may be greater, perhaps even by a factor of two, because we may not see the top of the prerift in the footwall. However, due to the lack of any deep borehole or shallow geophysical data (for the region between Kyparissi-Larymna; Fig. 6), values shown in Fig. 7 are the best available to date.

In Figs 6 and 7, pairs 1, 2, 3, 4, 5, 8, 9, 10, and 12 (see also Table 1) represent SubPelagonian zone pre-rift formations displaced across the fault, whereas pairs 6, 7, 11 and 13 are offset syn-rift rocks. Pairs 11, 12 and 13 yield real throw values ( + 10 m) because the hanging wall elevation of the base of the syn-rift is constrained from the geoelectrical data (profiles Ti and T, in Fig. 9). Pair 14 marks the end of the fault segment because at that locality, the synrift onlaps the fault-tip line. Separate curves have been constructed for the syn-rift and pre-rift throws. Figure. 7 shows that the throws associated with the syn- rift (thin dashed black lines) are at a maximum close to the centre of the fault segment we propose, decreasing towards the segment boundaries. This is also true for the throws associated with the pre-rift (Fig. 7, thin continuous black lines). Thus, the throw data

470 A. Ganas (‘1 trl

Table I. Mimmum throw estimates along the Atalanti Fault Segment. A set of 14 pairs of porn& arc used to

calculate throw variation from the aoutheastcrn tip (Larymna) to lhe northwestern tip (Atalanti) of the segment.

Length (in km) refers to a position of the intersection of a straight lint connecting each pair. with the fault tract,

and is measured from the southernmost origin (the SE trp). FW and HW refer to pomt elevations in metres acres,

the fault plane. respectively. The fourth column gives the minimum height difference between footwall and

hangingwall. The fifth and sixth columns provide the names of localities from the I:50 000 HAGS maps. and the

seventh column indicates the litho-stratigraphic identity of the pairs (after IGME, 1965). The hangingwall points

with negative elevations (depths below sea level) have been obtained from a geoelectrical survey and therefore

throws calculated there are actual rather than minimum values

Patr Length

I 0.0

2 3.9

7 6.0

-I 7.4

5 10.0

0 12.3

1 I.<.< s IS 7

0 70 0

IO 22.2

I I 24. I

I2 27.0

I3 32.0

I3 i5.0

FW HW 1 ‘hrou FW Locality

460 451

41s 359

320 120

366 I20

336 160

408 242

4ns 11:

530 so

603 I28

573 10

500 ioo

400 ilw

200

I30 140

56

200 246

176

I6h 272

550

175

53.3

x00

9x0

550

Stithos

Kapetamos

Scarp

Scarp

Martin0 W

Mytikas

L’oulgara

Vlilchog1or~o\

1.out\a

Kachr l-,)111\;1

Kolaka

(‘tilomon F,

HW Locality

Kokkinh

Koutroula

Shed E

Shed

Korsia I

Proskyna\

Trag. Beach

Tragana

Gaidouronisi

C&electric

Geoelectric

Geoelectrtc

Karagkro7i

Lithology

Jurassic (Lower)

Triassic Dolomite

Triassic Dolomite

Jurassic (Lower)

Jurassic (Lowet-)

Neogene

Neogenc

Peridotite

Jurassic (L-M)

Jurassic (L-M)

Neogene-base

Contact Triassic Dot,

Triassic Volcanics

Contact Neogene

Triassic Volcanica

Ncogene

support our imagery interpretation of the length and position of the Atalanti Fault Segment

and its segment boundaries. An independent confirmation of the throw values comes from Pair 15 shown on the map

in Fig. 6. The locality in the hangingwall (in Fig. 6: square Hl5, 2.5 km away from the fault) represents the base of the synrift. resting unconformably upon Jurassic limestones. The contact dips at about 17 degrees to the south. towards the Atalanti fault. By simple trigonometry (Fig. 6, inset), the depth of the synrift is projected 750 meters down-dip. of which 400 meters are the Neogene lacustrine formations (IGME, 1978). Similar rocks

(marls) were mapped by AG at an elevation of 304 meters by the Agia Trias Monastery south of the town of Tragana (in Fig. 6: square El 5). No basal unconformity was mapped at this locality. No other outcrops of these rocks occur at higher altitudes. However, we cannot safely conclude that the specific occurrence represents the top of the lacustrine part of the syn-rift at the footwall. Further southeast. in the Chiliadou region (Fig. 6) similar syn-rift rocks occur at an elevation of 408 metres, whereas in the Kyparissi region (to the northwest; Fig. 8). the uplifted syn-rift outcrops at 220 metres. If we accept an uplift of 400 metres (AMSL) for the top of the lacustrine syn-rift. then we obtain a total minimum displacement of Xl0 metres for the Tragana region, compared to a minimum throw of about 500 metres estimated from Fig. 7. We suggest that this independent evidence supports our throw determinations in Fig. 7.

We lack precise chronological data on the oldest syn-rift rocks offset by the fault.


0 1 km

Gulf of Evia

Skala n Ataiantonisi

Gaidouronisi

0 Elevation contour ___-*-_ Stream Trace of the Atalantl

. . . . ..- Roads b%*+

1 Geoh&fcal profile Fault Segment from

* satellite imagery

Fig. 8. Map showing the trace of the Atalanti Fault Segment (shown as a thick, light grey line), roughly following

the 1OOm elevation contour. Also shown are the lines of sections for the geo-electric surveys (thick black lines),

conducted by IGME across the Atalanti plain. An interpretation of lines Ti, T2 and T3 is shown in Figure 9.

Also shown are urban areas (cross-hatched circles), motorways (dark grey lines) and drainage (dashed black lines).

Note that all drainage to the north of the town of Atalanti flows to the East.

However, if we adopt a 3 Ma age (Pliocene) for the initiation of fault movement, based on the stratigraphy of the offset synrift south of Proskynas (following Rondoyianni-Tsiam- baou, 1984; Mettos et al., 1992; Fig. 6) and a minimum displacement of 810 metres for Tragana, a slip-rate of 0.27 mm/yr is implied. This value would increase in the Asprorema embayment region to 0.4mm/year, based on a slip value of at least 1200 metres (see minimum throws at that area in Fig. 7).

473

r

c fepth(m)

Atalanti Fault

lepth(m) 0 1 km

.lW

.O

.-loo

. -200

.-300

-400

.-500

depth(m) m Catbonate pre-rift

m Volcanic pre-fin

Fig. 9. Cross-sections depicting the interprctatlons of three electwal survey hnes across the Atalantl plain (after

Memou, 1986, unpublished report). The crowsect1ons arc ortented north-south, and from top to bottom of the

page are situated closer to the town of Atalantl. Notice the absence of any faulting inside the syn-rift. to the south.

Vertical exaggeration x 4.

4. KINEMATIC DATA

Again, in order to test our hypothesis concerning the length and position of the Ataianti Fault Segment we studied the strain pattern deduced from fault-slip data. Lineation data (N = 394; Fig. 10) were collected at five localities along the strike of the fault segment (black circles in Fig. 6), both from the main fault plane and from mesoscopic fault populations in the immediate footwall of the main fault. These lineations (frictional wear striae and corrugations) form during co-seismic motion and their survival on the fault surface renders them a reliable marker for extracting the azimuth of past co-seismic slip vectors.

Lineation data from the main fault plane were collected at three localities, at Asprorema,


t

c-----J

MarlIna N=lOB

Y

Atpmmma N=42 Chllladw N.166 (IAOM)

Fig. 10. Stereonets presenting poles to fault planes, lineations and fault-slip directions recorded along the strike of the Atalanti Fault Segment (from NW to SE). The ground distance between the localities Atalanti (NW) to

Martin0 (SE) is about 30 km.

Kyparissi and Martinon. Data from smaller footwall faults were collected south of the city of Atalanti, and in the Chiliadou quarry. Note that the dominant fault-slip direction in our data set does not remain constant along strike, but varies systematically. Fault-slip directions trend to the NE at the NW end of the fault segment defined above, whereas they trend to the NW at the SE end. Dominant slip-directions trend close to north in all other localities. Similar converging lineation patterns have been recorded along other normal fault segments of comparable length within the Gulf of Corinth rift system (Roberts, 1996) where a 90 degrees difference in mean trend was interpreted to mark the position of PSBs along this rift system. Thus, following Roberts (1996), this spatial variation in slip-direction supports our interpretation of the length and position of the Atalanti Fault Segment described above.

An interesting additional point is that two sets of lineations exist on the main fault plane at Martinon (see upper right stereonet at Fig. 10). The first set is characterised by major corrugations which, on average, plunge at 50 degrees towards N320 (NW). The second set is characterised by steeper, finer frictional wear striae on the corrugation surfaces, plunging 70 degrees towards NO28 (NE). We investigate the significance of these observations later in the discussion.

4.1. Summary of the length of the Atalanti Fault Segment and the positions of segment

boundaries Data concerning the footwall geomorphology, fault throws and fault-slip directions all

suggest that the Atalanti Fault Segment is about 34 km long, with the segment boundaries located to the northwest of the area shown in Fig. 4, and on the southeastern part of the Malesina peninsula, shown in Fig. 2, respectively. In Section 5 of the paper we examine

414 A. Ginas (‘I l/I.

whether the segment boundaries influence the extent of ruptures and whether the Atalanti

Fault Segment has grown through repetition of long (34 km). segment-boundary-to-segment boundary ruptures. or shorter (c. X 3) km) overlapping ruptures.

5. THE 1894 ATALANTI EARTHQl!c\KES: :\ I+‘IbLD RE-EXAMINATION OF THE

77:4. IX’)4 RUPTURE EXTENT .4ND ITS C‘OMPARISON TO THE F.4ULT SEGMENT

[LENGTH

In April I XY4 two dcv;astating earthquakes occurred in the region of Lokris (along the

southern coast of the Gulf of Evia). within a week (Skouphos. 1894; Mitsopoulos. 1895). Severe loss of life and property were reported along the coastal region. as well as surface

rupturing and ground-shaking phenomena (Fig. I I ). The severely damaged region was aligned in a general NW-SE direction with minor damage in the neighbouring areas of

Viotia (Fig. I) and the island of Ev/ia. Sea inundation was reported along the whole of the

23 00 E t 38 48’ N

EVIA ISLAND N

0 16km 1 I KOPAIS PLAIN

y 271411894 surface break (Ambraseys and Jackson, 1990)

. 1894 Population Centre

1894 surface break (confirmed during this study) : 1894 coastal subsidence

Fig. 1 I. Map showing the extent of the melsosersmal area on land for the 27 April 1894 earthquake (M 7.0)

according to Ambraseys and Jackson (1990), based primarily on the field report of Skouphos (1894).


coastline south of Livanates down to Almyra (Fig. 8), but to a lesser extent in the areas of Theologos (Fig. 6) and Larymna (Fig. 11). The earthquakes were attributed to slip along the Lokris Fault (Skouphos, 1894; IGME, 1989).

Skouphos, 1894 (pages 410-421) describes earthquake-related ground phenomena such as 100 m long ground-cracks with little or no throw resulting from the first event (20/4/1894) and shows that they were restricted to the Malesina Peninsula area (Fig. 11). In contrast, Skouphos says that damage associated with the second, stronger event (at 9:17 pm of 27/4/1894) extended for more than 30 km to the northwest, and major ruptures exhibiting both opening and throw were noted. Up to 2m of throw were recorded where the rupture crossed alluvium, with only 30 cm recorded associated with limestone bedrock. Skouphos suggested that, during the second earthquake, the major axis of the meisoseismal ellipse spanned a distance of 55 km in the NW-SE direction, a value that was interpreted by later workers (e.g. Lemeille et al., 1977; Stiros and Rondoyianni, 1985; IGME, 1989) to represent the co-seismic rupture extent. Although some short ground cracks (< lOO-150m) and landslides were mentioned by Skouphos in the region to the north and north-west of the town of Atalanti, no descriptions of throw values were included. Following re-examination of the geological and historical evidence, Ambraseys and Jackson (1990) suggested that a discontinuous rupture with 25 and 15 km long segments, on either side of the Atalanti town area, was produced by the second earthquake.

We have used (i) our own field mapping, (ii) structural and geomorphological interpretations of processed satellite data (Figs 13 and 14) and (iii) the recognition of long-lived segment boundaries in the areas northwest of the town of Atalanti and south of the town of Larymna (see previous section), to re-analyse the rupture extent. First we describe where evidence of ground rupture due to sub-surface fault motions is clear; we suggest that these areas are resticted to the region south-east of Atalanti (Fig. 6). Second, we describe our observations of the area north-west of Atalanti, up to the town of Agios Constantinos, where we suggest that the evidence of rupture related to sub-surface fault motions is less clear or lacking. We believe we have found preserved surface ruptures in the area between Martinon and Larymna (Fig. 11). Elsewhere we have reconstructed the rupture path from historical accounts (Skouphos, 1894; Mitsopoulos, 1895). We also use satellite observations, specifically that of shadows cast by faults and the rapid rise of morphological slopes (Figs 2 and 4) which occur along nearly straight lines in map view, to help us locate the likely positions of the rupture.

In the area 2 km south of Larymna, we mapped a single bedrock scarp, about a 1.5- l.Om high, for a NW-SE distance of more than 2 km. The rupture is marked a lO-30cm gap at its base, but shows no signs of ‘striping’ (at least detectable by eye) produced by previous earthquakes along its surface. The fault scarp extends up to an area 1.5 km southeast of the village of Martinon (Fig. 1 l), where a 2 m bedrock scarp bearing frictional wear striae, has been exhumed due to road building activity (present height 4m; Fig. 12a). This scarp recorded the 1894 break, which ranges between 60 and 30 cm in vertical displacement, as evidenced by the varying width of the 1st fresh stripe on bedrock above the vegetation line, and from the offset of the pre-1894 vegetation line itself. In the same Martinon locality, two higher, less fresh and perhaps older bedrock scarps are visible in the immediate footwall. These scarps show comparable heights to the one marking the position of the 1894 break, thus, creating a characteristic ‘stair-case’ footwall geometry (Fig. 12b). A similar geometry has been recognised in the Kaparelli region (Gulf of Corinth; Roberts et al., 1993) and may indicate recurrence of similar slip events, on the order of half a metre

Fig. 12. Field photographs of the Late Quaternary-Holocene bedrock scarps southeast of the village of Martinon

(April 1995). Both photographs are oriented SE-NW (from left to rtght). Top) the 1894 scarp in hmestone, recently

exhumed by road building. The height of the scarp is 4 metres. Arrows point to major corrugation crests along

the scarp. View towards the NW. Bottom) Two parallel bedrock scarps. of equal heights, forming a stair-case

geometry. Thick arrows point to the upper scarp. Scarp shown in the top photograph is enclosed in the square.

Thin arrows show the 1894 vegetation line. This locality is located by the arrow AF in the centre right of Figure 2. View is towards the SW. Horizontal distance from left to right is approximately 100 metres.


to one metre per event. Furthermore, 200m away on the adjacent hill to the southeast, a narrow opening at the base of the scar-p (10-20 cm wide) is partially filled with colluvium, leaves and other organic material. This opening is also suggested to be indicative of a recent surface break.

Towards the central part of the segment (Tragana area) no remains of the 1894 breaks have been found. However, there are two pieces of evidence for significant dip-slip movements due to large earthquakes: decimetre-size footwall uplift and historical reports. Foot- wall uplift may have occurred in the Tragana karstic springs area (symbol AF in Fig. 2) evidenced by an uplifted limestone notch. Field investigation by AG in October 1996 showed that the notch was created by fresh water erosion; no evidence of marine erosion or lithophaga borings were found. The notch can be traced at the same height above sea- level for a distance of at least 5 m, and is 6-10 cm wide. The notch exists 53 cm above a fresh-water spring. Thus, we interpret amount of uplift to be about 53 cm, perhaps produced by a single episode of rapid coseismic uplift. At a nearby locality, a narrow stripe, formed of algal material and interpreted to form between the high water and low water levels for spring discharge, is found 27cm above the present-day high discharge water-level. The stripe can be traced for a distance of greater than 3 m. Again, we interpret these observations to result from a single episode of rapid coseismic uplift. As there are no bedrock scarps to the north of the spring-line it is reasonable to assume that the 1894 rupture caused the observed uplift in these localities.

Also, in the Almyra region (between Tragana and Kyparissi; Fig. S), historical reports (see Mouyiaris, 1988 for summary) refer to co-seismic subsidence in classical times (426 BC), a phenomenon repeated some 2500 years later during the second 1894 earthquake (Skouphos, 1894; page 438). This supports the ideas presented above that metre-scale coseismic uplift/subsidence occurred in 1894.

Further north-west, where we have not found preserved ruptures, Skouphos describes surface breaks along the foothills of the Chlomon mountain (Fig. 5), as well as inside of, and in the vicinity of the town of Atalanti. Mitsopoulos (1895) in his field report refers to surface breaks, which may be the same as those described by Skouphos, extending from Atalanti to the Almyra region ‘in parallel sets to the main ground rupture’. Mitsopoulos (1895) also refers explicitly to the main ground rupture as starting from the Karagiozi channel (Ka symbol in Fig. 13) and continuing to the south-east. These observations indicate that the rupture mostly followed the prerift-synrift contact, that is the main fault plane. However, we did not find these ruptures, probably because of the friable nature of the faulted material (mostly spillites, greywackes and fanglomerates) and, most importantly, due to extensive agricultural activities in the area.

For the Kyparissi area (Fig. 8) Skouphos, 1894 (page 446) reported two surface breaks each with a 2545 cm aperture, enclosing an ‘semi-elliptical’ region (about 2 km long and 800 m wide). Local people claim that these 1894 surface breaks were still visible in the 1930s in contact with bedrock (‘Kokkinovrachos’ or red-rock locality), but we did not find them during our field examination. We did find small bedrock scarps (2-10 metres high, 200- 600 m long), aligned with the overall direction of the fault in the hilly area south and west of Kyparissi (Fig. 8), but they are not continuous, nor fresh. We also found one locality where a sharp offset of the vegetation line occurs across a bedrock fault scarp in the Asprorema region (Fig. 5b). However, in both these localities, we cannot be confident that these observations can be used to reconstruct the 27/4/l 894 rupture path; this needs further work. No field evidence was found for the 7 km crack observed by Skouphos in the region

478 A. Cianaa 1’1 III.

0 5km

Fig. 13. TM satellite image (1st prmcipal component) ol the arw 10 the northwest of the town of Atalantl. Small

black arrows point to the main population centres of the area: 41 IS Atalanti, Me is Megaplatanos (Skenderaga),

G is Couleml, Ze is Zeli, Lo IS Loggos and Ap is Agios (‘onstanbnos. respectively. Thick black arrows show the

postulated direction of the 27!4:1894 rupture (after IGME. 19X9). If that had crossed the Atalanti segment

boundary area. Symbol Ka points to the Karagiozi channel across where Mitsopoulos (1895) mentloned the start

of the main rupture developed to the southeast. Image is oriented 9 degrees east of North.

north and northwest of Kyparissi (ENE striking Moulki rupture, Skouphos, 1894, page

446), and later interpreted as a fault by Rondoyianni (1988). Also. no evidence for the so- called Moulki fault is present in the subsurftlce profiles of the syn-rift (Fig. 9; Memou,

1986). The region north-west of Atalanti is suggested to have been ruptured in 27i41 I894 (e.g.

IGME, 1989; Ambraseys and Jackson, 1990). However. our field-mapping and analysis of TM imagery has not provided evidence to support the existence of a 15 km continuous scarp suggested by IGME (1989). There is clearly no NW-SE fault-line visible in the principal component TM image of the region between Atalanti and Agios Constantinos (Fig. 13; symbols At and Ag). On the contrary, a drainage map (Fig. 14) extracted from the TM image shows that drainage west of Goulemi shows a trellis pattern and flows towards the north through a deeply incised gorge. These characteristics together with the development of the Loggos fan delta (Fig. 13, symbol Lo) point to the existence of long- lived drainage features which do not show any effects of recent disturbance by faulting. In addition, within the area to the east of Goulemi, drainage flows to the east. that is, it is


Populat Ion Cent re

.- l * catchment boundaries I’- stream / Hlgher-order stream

Fig. 14. Map showing a drainage interpretation of Fig. 13. Streams (continuous black lines) and the main watersheds (thick dashed lines) are indicated. Drainage south of the Gulf of Evia coastline up to the first E-W divide is not shown as it develops in the immediate footwall of the Kammena Vourla-Arkitsa Fault Segment (see Fig. 1 for location). NW-directed rectangle shows the hypothetical zone of ground ruptures postulated by IGME (1989). This drainage pattern is interpreted as resulting from long-term uplift of the footwall area of the Kammena Vourla Fault Segment. Note that the hypothetical rupture path crosses high mountains (see spot heights in black),

a feature that is inconsistent with ruptures occuring along pre-existing faults.

controlled by footwall uplift of the coastal normal fault segment (E-W ridges casting shadows in the top of Fig. 13). Note that within the (2 x 17 km) rectangular area shown in Fig. 14, where the 27/4/ 1894 ruptures are proposed to have occurred (IGME, 1989) drainage does not show any systematic offset that could indicate a recent disturbance.

This region did suffer ground-cracks resulting from the second earthquake (Skouphos, 1894, pages 446-447), but, due to the lack of a through-going fault we prefer to interpret these as the result of surficial movements due to (i) shaking; (ii) landslides; (iii) possible aftershock activity at the end of the rupture (28/4 to 30/4/1894; Skouphos, 1894, pages

459460) and (iv) minor slip on favourably oriented small (blind) faults (E-W to NW--SE). Thus. any ruptures in this region may not be the direct surface expression of fault motion at depth produced by the mainshock. We also suggest that coastal subsidence in the town of Agios Constantinos and the neighbouring Loggos area (Fig. 11) may have been due to

the liquefaction of unconsolidated coastal sediments. The coastal areas from Kammena Vourla to Arkitsa are part of the hangingwall of another, isolated normal fault segment

(Fig. I), the Kammena Vourla Fault Segment (Ganas et nl., 1996), which has developed independently from the Atalanti Fault Segment. and as such its footwall area (i.e. the region

Goulemi-Livanates) is unlikely to have been ‘invaded’ from ruptures from the latter. In summary. the hypothesis that the 27/4/1894 rupture propagated into the Renginio-

Kallidromon basin to the north. has not found support from our studies because: (i) no

north-west-striking fault scarp or surface break has been imaged from high-resolution (30m) satellite sensors (Fig. 13) or found during field mapping; (ii) the low relief in the

Atalanti segment boundary area (HAGS, 1975) indicates the position of a segment boundary that appears to have arrested rupture propagation throughout Late Quaternary times; (iii) Mitsopoulos (1895) refers explicitly to the main ground rupture as starting from the

Karagiozi channel (Ka symbol in Fig. 13). and (iv) the existence of undisturbed, eastward- directed drainage in the region to the north and bvest of the town of Atalanti (Fig. 14).

Also. by comparison with other normal faulting surface ruptures worldwide (Jackson and

White, 1989, Crone and Haller. 109 1). a rupture with only I -2 metres of throw in alluvium is likely to be significantly shorter than 55 km in length.

6. DISC‘USSION ~~ON(‘L.lJSIONS

The April 1894 Lokris earthquake sequence involved two large events (M 2 6.0) within

a week’s time, along NW~-SE striking large. planar normal faults. in an actively extending

continental area of average crustal thickness (3 2 36 km: Makris, 1985). Our field obser-

vations, reviews of the reports of Skouphos ( 1894) and Mitsopoulos ( 1895), and interpret-

ations of processed, high resolution satellite imagery. suggest that in 18Y4, the entire length of the Atalanti Fault Segment was ruptured, that is. a distance of about 34 km.

However. we do not know if(i) the ruptures result solely from the second earthquake.

or (ii) from a combination of motions produced by both earthquakes. If the first hypothesis is correct, the 27/4/l 894 earthquake comprised a segment boundary-to-boundary event, in the sense of the fault segment definition suggested by Roberts and Koukouvelas (1996). This would be the longest rupture from a single earthquake recorded in central Greece. Also. if this 27/4/1894 rupture length is accepted. then it is reasonable to interpret the 2Oi4/1894 (M6.4; Ambraseys and Jackson. 1990) event as rupturing a neighbouring fault segment to the south, along the Gulf of Evia. If we also accept a northwestward rupture propagation for the first event, then the large amount of destruction in the Malesina peninsula may be also explained. This area is part of the Larymna segment boundary where the 20/4/1894 rupture may have terminated, thus. it may have received most of the radiated seismic energy. Perhaps the first earthquake triggered the second because, following the lead of other workers who suggest that stress-levels on faults can be affected by rupture on neighbouring faults (Stein et ul., 1992; Hubert ct al.. 1996; Hodgkinson e’t al., 1996), a rupture to the south-east may have loaded the Atalanti Fault Segment, and perhaps brought it closer to failure. The effect is enhanced because it is probable that the 426 BC Gulf of Evia earthquake (e.g. Mouyiaris, 1988) may have ruptured the Atalanti Fault Segment; the


elapsed time of 2500 years may be great enough so that the segment was near the end of the quiescent stage of its seismic cycle. Alternatively, the 20/4/1894 event may be due to an offshore active fault, as Skouphos, 1894 (page 461) suggested, himself. Moreover, the first rupture hypothesis implies that unusually long ruptures (> 20 km) along other fault segments in Central Greece are not to be ruled out. Most of these fault segments are 30-35 km long (e.g. Roberts and Koukouvelas, 1996), and as such they may be able to produce exceptionally long ruptures. In turn, such rupture lengths may be associated with events greater than those currently expected (Mmax=6.8-7.0; IGME, 1989).

The second hypothesis involves partial rupture of the Atalanti Fault Segment. The 20/4/1894 earthquake may have ruptured the south-eastern part of the fault segment because Skouphos describes localised damaged and ground cracks in this area following the first event. Then, the 27/4/1894 earthquake would have ruptured only the north-western part, from Tragana to Atalanti. This implies 10-15 km rupture lengths for each event, a length that is similar to better constrained events elsewhere in Greece (Ambraseys and Jackson, 1990, Pavlides et al., 1995; Roberts and Koukouvelas, 1996). A triggering-effect scenario could still apply, and the hypothesis is similar to that of the 1981 Alkyonides earthquakes where two (> MS 6.0) events occurred within a few hours on the same fault segment (compare Jackson et al., 1982; Roberts, 1996, Hubert et al., 1996). Unlike the first hypothesis, where characteristic behaviour (Schwartz and Coppersmith, 1984) and perhaps a uniform recurrence interval is possible, the second hypothesis cannot involve uniform recurrence intervals. This is because Roberts (1996) has shown that if only short portions of longer faults are ruptured during single earthquakes, then geometric considerations demand that recurrence intervals must vary along faults and with time at single localities, Thus, a probabilistic approach to hazard analysis using existing techniques is perhaps valid with the first hypothesis, but not for the second.

Palaeoseismological trenching studies may help to decide if the Atalanti Fault Segment exhibits behaviour associated with segment boundary to segment boundary events or events that are shorter than the segment. However, an alternative approach may be to examine lineation data sets as suggested by Roberts (1996). Earlier we noted that two sets of lineations exist on the main fault plane at Martinon (Fig. lo), the older set plunging towards N320 (NW), and the younger set towards NO28 (NE). If these striations were produced by ruptures with converging patterns of slip towards the hangingwall, as is the case for normal faulting ruptures (e.g. Roberts and Koukouvelas, 1996), then two sets with different orientations implies rupture of different portions of the same fault segment during successive earthquakes. This rather indirect evidence supports the second hypothesis. However, we must re-call that in the account of Skouphos (1894) no mention of ruptures associated with the first event are noted, and this supports the first hypothesis.

The growth of seismogenic normal faults is of great interest because earthquake segmentation may be influenced by growth patterns (e.g., Cowie and Scholz, 1992; Roberts, 1996; Main, 1996). The data sets we collected for the Atalanti Fault Segment permits us to formulate a linkage-growth model for this fault. We note the left-stepping nature of the topographic contours in the region to the southeast of Atalanti (Fig. 6; HAGS, 1975) and suggest that the formation of the Asprorema embayment (Figs 4 and 8) is probably due to the linkage of two, overlapping (in the extension direction), Atalanti fault proto-segments, located either side of the Kyparissi limestone hills region (Fig. 15; south and west of Ky mark in Fig. 4). The limestone ridges area (together with the thin cover of lacustrine syn- rift) to the west of Kyparissi before linkage would be part of the hangingwall of the western

Gulf of Evia

0 5 IOkm

Persistent Segment Boundary Extinct Segment Boundary

\

Trace of Atelanti Fault Segment l Population Centre Protosegment from sateflite imagery

Martinnn area shows the southeasterl! dlrcction 01’ fault growth after linkage.

proto-segment. and after linkage It was transferred to the Atalanti Fault footwall and uplifted to its present elevation (200 m; Fig. X and 15). Also. before linkage, the area to the

east of Kyparissi would be located at the hangingwall of the eastern proto-segment and probably at its termination, It follows that the area around Kyparissi before linkage may have been characterised by low relief. and perhaps by transverse bedrock ridges

development, i.e. it may comprised an ancient segment boundary. that is now extinct. This model explains both the low footwall elevation and the high hangingwall elevation

of the pre-rift (e.g. the bedrock ridge of Gaidouronisi; IGME, 1965) observed today in the Asprorema-Kyparissi area. However, in the throw distribution profile (Fig. 7). the fault throws between kilometres 22 and 24 do not decrease because of this linkage; instead, the amount shown is the summed throw across the fault (as this is now a single kinematic entity), so that throw values remain comparable with the regions both to the SE and to the NW.

This fault growth model is supported by the similar morphometric characteristics of the footwall basins (‘wine-glass’ valleys) south of Tragana and south of Atalanti, respectively (Figs 3 and 4). Such equal-in-size erosional features may have initiated at about the same


time (3 My-age of the oldest offset synrift), and developed ‘in phase’ and orthogonally to the two parallel proto-segments. According to the same model, the two bends seen in Fig. 2 (symbol B) are not independent, smaller segments as suggested by Poulimenos and Doutsos (1996) but they are related to pre-faulting irregularities, linked together during the growth of the eastern proto-segment. It is also suggested that, segment linkage must have happened relatively late in the fault history (Mid-Upper Pleistocene?), because of the abandonment of three, ‘wine-glass’ valleys (Fig. 4; valleys numbered 8 to 10) that may have been located in the overlap region. These valleys are of comparable morphometry and orientation to those on the Chlomon Mountain further north-west (Fig. 5a) and were part of the footwall drainage (the hinterland) of the western proto-segment. After linkage, because of the uplift of the area west of Kyparissi these three streams changed their flows towards the north-west to join with that of valley 7 in the Asprorema region. A late linkage is also suggested by the recent exhumation of a large, uneroded fault plane running parallel to the Asprorema stream (Fig. 5b), that is not seen to propagate further to the west.

It is interesting to point out that the 27/4/l 894 rupture hypothesis (34 km length) requires the rupture path to go through the linkage (Asprorema) area, and no surface ruptures to occur in the newly-uplifted area to the south of it. The extent of the 1894 ruptures as reported by Skouphos (1894) and Mitsopoulos (1895) was indeed confined within the Atalanti plain and in the Almyra region (Fig. 8). Equally important, if the second hypothesis is true then, the area where the 1894 ruptures terminated (or started) is the Tragana- Kyparissi (Almyra) area, i.e. the extinct segment boundary (Fig. 15). This again is consistent with the field observations of Skouphos and Mitsopoulos, who reported ground ruptures and sea inundation in that area.

This study emphasises the importance of combined data sets in identifying the lengths of faults: in this case satellite imagery interpretations, topographic data, throw data and kinematic data. We suggest that such studies supply segment length estimations that are more robust than studies using only map traces of faults to define segment lengths. Furthermore, our study shows that it is important to identify both the length of the fault segment and the rupture length during study of normal fault earthquakes. Only then can the long-term behaviour of the fault segment (including numerous seismic cycles), be envisaged: this has dear implications for assessing the hazards associated with large normal- slip earthquakes.

Acknowledgements-We are grateful to the IGME Laboratory for Remote Sensing (Athens) for kindly providing AG with Landsat data, and IGME International Relations for field work permits to GPR. The Hellenic State Scholarships Foundation (IKY) supported this research with a doctoral studentship to AG. GPR was funded by NERC (GR9/1034) and Birkbeck College. We wish to thank Prof. Geoff Wadge (ESSC, Reading) and Dr. Kevin White (Geography, Reading) for many discussions on the use of remote sensing in geological studies. Iain Stewart pointed us to the Tragana springs locality. We are also indebted to Vaggelis Spyridonos for translating the Skouphos 1894 paper and to Victoria Buck for discussions regarding the 426 BC earthquake and making available to us related literature. Yiannis Bacopoulos, Steve Thorne and Louise Seymour assisted in the field. Heather Browning drew Figures 1 and 9. We thank Iain Stewart for a helpful review and three anonymous referees.

REFERENCES

Agar, S. M. and Klitgord, K. D. (1995) Rift flank segmentation, basin initiation and propagation: a neotectonic example from lake Baikal. J. Geological Society, London 152, 849-860.

Ambraseys, N. N. and Jackson, J. A. (1990) Seismicity and associ.ated strain of central Greece between 1890 and 1988. Geophys. J. Int. 101,663-708.

484 A. C;anas t11 trl

Anders, M. H. and Schlische, R. W. (1994) Overlapping faults, intrabasin highs, and the growth of normal faults. J. GeologJl 102, 165180.

Armijo, R., Meyer, B., King, G. C. P., Rigo, A. and Papanastasiou, D. (1996) Quaternary evolution of the Corinth Rift and its implications for the Late Cenozoic evolution of the Aegean. Geophys. J. Znt. 126, I l-53.

Billiris, H., Paradissis, D., Veis. G., England, P.. Featherstone, W., Parsons, B.. Cross, P., Rands, P., Rayson. M., Sellers. P.. Ashkenazi. V., Davison, M., Jackson, J. and Ambraseys. N. (1991) Geodetic determination of tectonic deformation in central Greece from 1900 to 1988. Nature 350, 124.-129.

Bushara, M. N. (1996) Neogene geologic structures using enhanced Thematic Mapper over eastern Zagros Mountains. Int. J. Remote Sensing 17(2), 303-3 13.

Clarke, P. J., Paradissis, D., Briole. P.. England, P. C., Parsons, B. E., Billiris, H., Veis, G. and Ruegg, J.-C. (1997) Geodetic investigation of the 13 May 1995 Kozani-Grevena (Greece) earthquake. Ge0phqJ.s. Re.v. Letters 24, 707~ 7 10.

Cowie, P. A. and Scholz, C. H. (1992) Growth of faults by accumulation of seismic slip. J. Geophys. Reseurch 97, 11085511095.

Crone. A. J. and Haller. K. M. (I 991) Segmentation and the coseismic behaviour of Basin and Range normal faults: examples from east-central Idaho and southwestern Montana, U.S.A.. J. Struct. Geobgy 13, 151 164.

Ganas, A. and White, K. (1996) Neotectonic fault segments and footwall geomorphology in Eastern Central Greece from Landsat TM data. Geol. Sot. Greece Sp. Pub/. 6, 169- 175.

Ganas, A., Wadge, G. and White. K. (1996) Fault Se~gmentation and Tectonic Geo- morphology in Eastern Central Greecefiom Satellite Data. Eleventh ERIM Conference on Applied Geologic Remote Sensing. 27 29 February 1996, Vol. I. pp. I 19- 128. Las Vegas.

HAGS (I 975) (Hellenic Army Geographical Service). lstiaia Map Sheet, I : 100000. Athens. Hodgkinson, K. M., Stein, R. and King, G. C. P. (1996) The 1954 Rainbow Mountain-

Fairview Peak-Dixie Valley earthquakes: a triggered normal faulting sequence. J. Ge0phy.s. Res. 101, 25459 -2547 1.

Hubert, A., King, G.. Armijo. R., Meyer, B. and Papanastasiou, D. (1996) Fault re- activation, stress interaction and rupture propagation of the 1981 Corinth earthquake sequence. Earth Plan. Sci. Letters 142, 573 585.

IGME (1965) Geologicul mup of’Greecr. Sheet Atalanti. Scale 1:50000, Athens. IGME (1978) Geological map of’Gr_cece. Sheet Larymna, Scale 1:50000. Athens. IGME ( 1989) Sei.snlotcc,tortic. mup c$Grww. Scale I :500000. Athens. Jackson, J. A., Gagnepain, J., Houseman. Cr., King, G. C. P., Papadimitriou, P.. Soufleris.

C. and Virieux, J. (1982) Seismicity. normal faulting, and the geomorphological development of the Gulf of Corinth (Greece): the Corinth earthquakes of February and March 198 1. Earth Plan. Sci. Letters 57, 377- 397.

Jackson, J. A. and White, N. J. (1989) Normal faulting in the upper continental crust: observations from regions of active extension. J. Struct. Geologlx 11, 15- 36.

Jackson, J. and Leeder, M. (1994) Drainage systems and the development of normal faults: an example from Pleasant Valley, Nevada. J. Struct. Geology 16, 1041-1059.

Karnik, V. (197 1) Seismicit?, of‘the Europeon urea. Part 2. 2 I8 pp. D. Reidel Publishing Co, Dordrecht-Holland.

Leeder. M. R. and Jackson. J. A. (19Y3) The interaction between normal faulting and drainage in active extensional basins, with examples from the western United States and central Greece. Basin Res. 5, 79- 102.

Lemeille, F.. Gauthier, A.-J., Jarrige, J.-J. and Philip. H. (1977) Evolution neotectonique du domaine egeen et de sa bordure externe orientale: une mise au point. BUN. Sot. GPol. France X1X(7), 673 -677.


Main, I. (1996) Statistical physics, Seismogenesis and Seismic Hazard. Rev. Geophysics 34(4), 433462.

Makris, J. (1985) Geophysics and Geodynamic Implications for the Evolution of the Hellenides. In Geological Evolution of the Mediterranean Basin, eds. D. Stanley and F. Wezel, pp. 231-248. Springer.

Mettos, A., Rondoyianni, Th., Ioakim, Ch. and Papadakis, I. (1992) Evolution gko- dynamique et reconstruction palCoenvironmentale des basins nCog&es-quaternaires de la G&e centrale. Paleontologia I evolucid 24-25, 393402.

Meyer, B., Armijo, R., Massonnet, D., de Chabalier, J. B., Delacourt, C., Ruegg, J. C., Achache, J., Briole, P. and Papanastasiou, D. (1996) The 1995 Grevena (Northern Greece) Earthquake: Fault Model constrained with tectonic observations and SAR inter- ferometry. Geophys. Res. Letters 23,2611-2680.

Memou, Tz. (1986) Geophysical research in the Atalanti area. IGME unpublished report, 24 p. (in Greek).

Mitsopoulos, K. (1895) The Lokris mega-earthquake, (in Greek), 40 p. Athens. Mouyiaris, N. K. (1988) Destructive historical earthquakes in N. Euboicos and Maliakos

Gulfs-Their significance to the evolution of the area. In Engineering geology of ancient works, monuments and historicalsites, eds. Marinos and Koukis, pp. 1249-1256. Balkema, Rotterdam.

Papazachos, B. and Papazachou, K. (1989) The earthquakes of Greece, 356 p. Ziti Editions, Thessaloniki.

Pavlides, S. B. (1993) Active faulting in multi-fractured seismogenic areas; examples from Greece. Z. Geomorph. N. F., Suppl.-Bd. 94, 51-12.

Pavlides, S., Zouros, N., Chatzipetros, A., Kostopoulos, D. and Mountrakis, D. (1995) The 13 May 1995 western Macedonia, Greece (Kozani-Grevena) earthquake; preliminary results. Terra Nova 7(5), 544-549.

Poulimenos, G. and Doutsos, T. (1996) Barriers on seismogenic faults in Central Greece. J. Geodynamics 22, 119-135.

Richter, C. F. (1958) Elementary Seismology, 768 p. W.H. Freeman and Co, San Fransisco. Roberts, G., Gawthorpe, R. and Stewart, I. (1993) Surface faulting within active normal

fault zones: examples from the Gulf of Corinth fault system, Central Greece. Z. Geomorph. N. F., Suppl.-Bd. 94, 303-328.

Roberts, G. P. and Gawthorpe, R. L. (1995) Strike variation in deformation and diagenesis along segmented normal faults: an example from the eastern Gulf of Corinth, Greece. In Hydrocarbon Habitat in Rift Basins, ed J. J. Lambiase. Geol. Sot. Sp. Publ. 80, 57-74.

Roberts, G. P. and Koukouvelas, I. (1996) Structural and seismological segmentation of the Gulf of Corinth Fault System: implications for models of fault growth. Annali di Geophysics XXXIX, 619-646.

Roberts, G. P. (1996) Noncharacteristic normal faulting surface ruptures from the Gulf of Corinth, Greece. J. Geophys. Res. 101,25255-25267.

Roberts, S. and Jackson, J. (1991) Active normal faulting in central Greece: an overview. In The Geometry of Normal Faults, eds. A. M. Roberts, G. Yielding and B. Freeman. Geol. Sot. Sp. Publ. 56, 125-142.

Rondoyianni-Tsiambaou, Th., (1984) Etude Neotectonique des rivages occidentaux du canal d’Atalanti (G&e Centrale), 200 p. Doctoral Dissertation, Universitl: de Paris Sud.

Rondoyanni, Th. (1988) Tectonique rCcente et subsidence des sites historiques aux rivages de Locride, G&e. In Geologie de l’ingenieur appliquPe aux travaux anciens, monuments et sites historiques, eds. Marinos and Koukis, pp. 1583-l 589. Balkema, Rotterdam.

Sabins, F. F. (1987) Remote Sensing. Principles and Interpretation, (Second Edition), 449 p. W.H. Freeman and Co., New York.

Salomonson, V. V. and Stuart, L. (1989) (Editors). Thematic Mapper Research in the Earth Sciences. Remote Sens. Environ. 28, 5-347.

486 A. Ganas 6’1 (~1.

Schwartz, D. P. and Coppersmith, K. J. (1984) Fault Behavior and Characteristic Earth- quakes: examples from the Wasatch and San Andreas fault zones. J. Geuphys. Research 89(B7), 5681-5698.

Skouphos, T. (1894) Die swei grossen Erdbeben in Lokris am X/20 und 15,/27 April 1894. Zeitschr$ Ges. Erdkunde x Berlin 24, 409 474.

Stein, R. S., King, G. C. P. and Lin, J. (1992) Change in Failure Stress on the southern San Andreas fault system caused by the 1992 magnitude = 7.4 Landers earthquake. Scienw 258, 1328-m 1332.

Stiros, S. C. and Rondoyianni, Th. (1985) Recent Vertical Movements across the Atalanti Fault-Zone (Central Greece). Pa,geoph. 123, 837- 848.

Wallace. R. E. (1978) Geometry and Rates of change of fault-generated range fronts. North-Central Nevada. J. Rc~srarch L:.S. Geol. Surzv?~ 6(5), 637-650.

Westaway, R. (1991) Continental extension on sets of parallel faults: observational evidence and theoretical models. In The Geometry of Normal Faults, eds. A. M. Roberts, G. Yielding and B. Freeman. Geol. So<.. Sp. Puhl. 56, I43 - 169.

Zhang, P., Slemmons, D. B. and Mao. F. (1991) Geometric pattern, rupture termination and fault segmentation of the Dixie Valley-Pleasant Valley active normal fault system. Nevada, U.S.A. J. Struct. Geolqq~~ 13, 165 176.

segment boundaries, the 1894 ruptures and strain …

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