Fractal analysis of normal faults in northwestern Aegean area, Greece

J. Geodynamics Vol. 26, No. 2-4, pp. 197-216, 1998

Pergamon 0 1998 Published by Elsevier Science Ltd

All rights reserved. Printed in Great Britain PII: SO264-3707(97)00052-5 026&3707/98 519.OOt 0.00

FRACTAL ANALYSIS OF NORMAL FAULTS IN NORTHWESTERN AEGEAN AREA, GREECE

THEODOR DOUTSOS* and IOANNIS KOUKOUVELAS

Department of Geology, University of Patras, 26101 Patras, Greece

(Received 18 November 1996; revised 3 June 1997; accepted 13 August 1997)

Abstract-Normal faults within the Ptolemais coal field and large seismogenic faults in the northwestern Aegean remain fractal for displacement values larger than about 1 m. The kinematic parameters on reverse drag profiles such as length of rollover, footwall uplift and wavelength of footwall uplift show that all three parameters have a power law relationship, expressed by a c exponent of about 1, with the maximum displacement which take place across the fault. Footwall uplift/hanging wall subsidence ratio is about l/2.

The displacement analysis help us to propose a growth model for larger seismogenic faults in the NW Aegean, as is the ‘Hepiros fault set’ and the ‘Aliakmon fault zone’. Faults within the ‘Aliakmon fault zone’ were independently developed, at the first stages of deformation, by tip line deformation and out-of plane bifurcation, whereas later, deformation continued by segment linkage. One of these faults the ‘Sarakina fault’ was react- ivated during the 1995 earthquake to produce a 25 km long surface rupture. A long term slip rate of about 0.3 mm a-’ has been estimated by taking into consideration that over the past 6 Ma a maximum displacement of 1700 m across this fault has taken place. 0 1998 Published by Elsevier Science Ltd. All rights reserved

INTRODUCTION

The displacement field of a blind normal fault has been described to have an elliptical shape with the fault surface as a principal plane (Rippon, 1985). Fault surface ellipses have an horizontal major axis termed as the ‘fault width’ (1) and a short axis ‘down-dip fault height’ (h) (Fig. 1) (Walsh and Watterson, 1989).

The maximum displacement (&ax) occurs at the central part of the fault surface and diminishes toward their edges (Fig. 1). In addition displacement decreases to zero along a normal to the fault surface to form the characteristic flexural profile of several normal faults (Gibson et al., 1989). The variation in displacement along this direction is expressed by a longer hanging wall rollover (r,) and a shorter wavelength of footwall uplift (r-J.

The relationships between these geometric and kinematic parameters have been recog-

*Author to whom all correspondence should be addressed: E-mail: [email protected]; Fax: (61) 994485.

197

198 T. Doutsos and I. Koukouvelas

n = wwzlmph of hangingwall rollover, rt = wweleqth of footwall uplin Max= maximum diit. I/z= faub half-width. h= dowdip fault hd8ht

Fig. I. Geometric parameters of a normal fault.

nized as fractal in nature in many faults (Walsh and Watterson, 1989; Marrett and

Allmendinger, 1991; Cowie and Scholz, 1992a). However, often the scatter in the diagrams of the collected parameters is considerable large so that standard statistical methods cannot

be applied (S. Gillespie et al., 1993). These departures from self-similarity of faulting are caused by host rock properties (Cowie and Scholz. 1992b). fault linkage (Cartwright et al.,

1995; Willemse et al., 1996) as well as by the different growth stages in the development of a fault population within a progressively strained area (Wojtal, 1996).

The Aegean region, one of the most rapidly extending continental areas in the world (McKenzie, 1978) is a typical area of analysis of normal faults, (Jackson et al.. 1982; King

et al., 1982; Doutsos and Poulimenos. 1992, Pavlides, 1993). But. no fractal techniques have been employed yet to establish the scalmg properties of these faults.

In this paper we first check the power-law hypothesis on neotectonic normal faults, whose

data span more than three orders of magnitude in fault size. Then we recognize faults with displacement > 1 m to be scale invariant and determine the parameters of the power law. Subsequently we show that geometry and kinematics of larger seismogenic faults in the northwestern Aegean is predictable and we present a growth model of faulting.

TECTONIC SETTING

Northwestern Aegean includes the northern part of the Hellenic mountain chain which was built up during the collision between the Apuliay plate and the Eurasian continent in the Middle Miocene (Dewey et al., 1973; Mountrakis, 1986; Doutsos et al.. 1993). Since the Upper Miocene, the whole crust has thinned and collapsed to form a complex pattern of intramountain basins and highs with the characteristic basin and range type topography

Fractal analysis of normal faults in northwestern Aegean area 199

(Fig. 2, inset) (Doutsos, 1980; Pavlides, 1985; Caputo and Pavlides, 1993, Doutsos et ul., 1994). The older syn-rift sediments of late Miocene-late Pliocene age were deposited within the Ptolemais and Kozani basins (Fig. 2) and comprise a basal conglomerate-sandstone unit that is followed by a fining upward sequence of lacustrine-fluvial sand-clays grading upward into lignite horizons (Schneider and Velitzelos, 1973; Ioakim, 1981; Koufos, 1982). This sequence reaches a thickness of more than 600 m and is capped by fluvial and alluvial mudstone and sandstone of Pleistocene age (Faugeres, 1966). Further west in NW Epiros province (Fig. 2) young basins contain a thin and very young sequence of conglomerates, breccia slope talus and red fluvial and alluvial deposits.

The coal field within the Ptolemais basin provides an excellent opportunity to analyze normal fault populations. The region provides several artificial outcrops (up to 2500m long), which allow sampling of a large number of faults with displacements ranging over several orders of magnitude. Another optimum situation within the coal field exposures is the fact that they comprise a single data set of normal faults whose direction change only slight from ENE-WSW to E-W (Fig. 3 nets l-5). Faults are planar, with dips of 60-80”, and show a near tip listric geometry.

The Ptolemais basin is a 60 km long and about 20 km wide subsidence area which trend parallel to the NW trending isopic zones of the Hellenides. As normal faults within the basin trend exclusively ENE-WSW (Fig. 3) it seams very probable that the northwest trending morphology of the basin follows an inherited bathymetry induced by a previous

wridimmtr

mbasomont

i i il

Fig. 2. Post Miocene basins in northwestern Greece. The inset shows the post-Miocene rift belts transversely

cross-cutting the Hellenic peninsula. Hatched box: study area, rift belts: VLG, Volvi-Lagada graben; AG,

Anthemouda graben; ALG, Almyros graben; SG, Sperhios graben; AMG, Amvrakia graben; TG, Trichonis

graben; EG, Evoikos graben; CG, Corinth graben; PTG, Patras graben; PG, Pyrgos graben; KG, Kalamta graben;

HT, Hellenic trench.

200 T. Doutsos and 1. Koukouvelas

,t.Jfifi 41

-. ___ _ - - --___

\ I LEGEND

Fig. 3. Simplified tectonic map of the Ptolemais basin(after Pavlides, 1985). I ~5 stereo-plots showing slickensides

on mesoscopic faults. CJI (maximum compressive stress). r2 (intermediate compressive stress) and 03 (minimum

compressive stress) are computer estimated.

late Miocene extensional phase (Doutsos EI ul., 1994). The basin is separated from the Vernon and Voras mountains to the north by a northeast trending fault zone consisting of

two west stepping segments: the Petres and Vegoritis segments (Fig. 3). Displacement along this fault zone reach a maximum of 1300m (see Fig. 5 in Pavlides and Mountrakis, 1987).

To the south the basin is bounded by the Drepano fault, a west trending and 15 km long fault, along which the basin floor is subsided about 400m (Koukouzas and Kotis, 1993). Internal deformation, within the basin sediments, is represented by a conjugate system of

nearly NE-SW trending normal faults (Fig. 3: diagrams l-5). Fault planes are recognizable as zones typically 1 mm- 1 m wide consisting of thin cataclasites and fault gouge. They are often striated and show variations in dip between 55 -70 The genera1 strike of these faults

changes from NE-SW in the northern part of the basin to E-W in the southern part. Stress analysis based on computer calculations reveals that these faults were formed by a tensional stress which change from a NW-SE direction to the north into a N-S direction to the south.

FRACTAL ANALYSIS

Cumuluti~le,frequen~~ plot &+uc~turr displacements

Ninety individual faults were measured by one dimensional sampling of displacements along three lines in the northern and southern Ptolemais coal field as well as in the Amynteon

201

i -I

1.0 4

0.0

Fractal analysis of normal faults in northwestern Aegean area

open triangle= 0.58 open diamonds= 0.46 crosses= 0.48 solid circles= 0.52 solid d&nods= 0.44

01 1.0 10.0 loo.0 lam.0 1oml.o

Swum)

Fig. 4. Cumulative frequency plot of fault displacements in the Ptolemais basin (for details see text).

coal field (Fig. 4). Often the trend of these sample lines is imposed by the nature of excavation and are oriented at a high angle to the strike of the faults. In that cases the sample line was reoriented to be perpendicular to the NE-SW trending faults by the application of simple trigonometric corrections. Displacements range from 5 mm to 100 m (Fig. 4) and have been measured in every fault across the three sample lines either in the field or from cross sections (Fig. 5) using the separation of marker beds. Only the maximum displacement on each fault was estimated. As sample lines are completely exposed the regional extension can be obtained by summing heaves on individual faults. Thus, estimates of total extension across the three sample lines range from 15% to 35%.

The displacement distribution curves of the three sample lines in the plot of Fig. 4 show a segmented geometry. The central part of each one curve defines a straight line, indicating a power-law relationship between fault displacement and cumulative number. The straight

I

B 18OOm /

Fig. 5. Reverse drag profiles in the southern (a) and in the northern (b) coal fields of the Ptolemais basin.

202 T. Doutsos and 1. Koukouvelas

lines extend over two orders of magnitude of displacement size and have slopes between 0.44-0.48. Small breaks in the slope of the curves are not artifact of the sampling techniques but represent abrupt changes in the lithology of the host rock. The left hand segments of the curves have shallow slopes suggesting that small faults may have been inadequately

sampled or that small faults, with displacements ~0.7 m, obey a different fault mechanism than the larger faults.

The steep right hand segment of the curves line (x) in Fig. 4, which represents faults with

displacements greater than approximately 70m. suggests that larger faults were under- sampled. This is very probable as the three sample lines are located in the internal parts of

the Ptolemais basin and large, basin bounding faults, were omitted and not statistically interpreted.

In order to avoid this error (censoring) we suppose that all faults in the sample area belong

to the hanging wall of a basin forming fault. like the Petres fault. The real displacement of 1100m of this fault, can be predicted. if the entire fault population curve (Fig. : line (u) extends to intersect the ordinate of (d) at I 1OOm. Although an observation gap of faults

with displacements between 100m and IOOOm exist it seams therefore very probably that all faults with displacements >0.7m are scale invariant. Thus statistical bias caused by

undersampling of larger faults can be omitted. if the steep dipping segment of the curve (x)

is replaced by a shallower dipping segment (u) which is the downward prolongation of the

central straight segment. Fractal dimension (D) of the entire fault population is 0.52, and lies within the lower end

of the range of 0.37 < 1.7 noted for other fault populations (Marrett and Allmendinger.

1991). This implies a greater proportion of large to small faults than in many faulted terrains. Small values of the fractal exponents have, in fact, been attributed to extensional strain in seismogenic areas with a h value between 1. I5- I .25 (Pickering et al., 1996), which

is clearly the case of our study area.

Displacement-dowx dip hri.yht plot of fractures Displacement and down dip height of 24 blind faults were measured directly in the field

or from photographs and have no significant errors. Although a large scatter in don h and vice versa can be shown in Fig. 6 (faults of similar down dip height can have displacements that differ by a factor of 2) the distribution of the data is consistent with two parallel lines

(I and -_) with a common slope of I. (S. Clark and Cox, 1996). Line ($) is derived from small faults with d < 0.4m and h < 5-7m and have a ratio

logdmaxllogh = 0.05, If we assume that the length (I) of the fault is equal to twice the

downdip height of the fault (h) (Barnett Pt al.. 1987) then we can calculate the scaling parameter 1’ (dmaxil sensu Cowie and Scholz, 1992a) for this relationship to be 0.015. This low value of ;’ is characteristic of rocks of low shear strength such the coal beds in the British coal fields (Walsh and Watterson. 1988) or the weak layered deposits of the basin and range province (Carter and Winter, 1995).

Line (z) is derived from larger faults with d between 0.4m-1.5 m and h between 5rn- 23 m. The scaling parameter 7 = 0.04 is larger than that of smaller faults caused very probably by different host rock properties. The larger faults cut not only coal beds but often they extend into a mixed sequence of coal beds and diagenetically compacted sandstones and mudstones, which have greater shear strength and higher frictional resistance than the coal beds alone.


0.10 , I 7 ,I/!.‘, I1l1l1

1.00 10.00 100.00 Log h (hrwidth &fault, in m)

Fig. 6. Displacement-down dip height plot of faults in the Ptolemais basin(for details see text).

As the vertical extent of the coal-field outcrops is limited, the y parameters can be representative only for faults with h up to 23 m.

Reverse drag projiles In order to study the self-similarity relationships of faults with length > 46 m (down dip

height 23 m)detailed displacement relationships profiles were constructed parallel to the slip vector of the faults. They extend up to 2000m and show the decrease of penetrative deformation away from the faults (Fig. 5). Only single isolated faults are considered and the marker horizons are at the same elevation at both ends of the deformed region (Figs 1 and 5). If I> r > h and I = 2h an approximation which observation and theory show to be valid (Savage and Hastie, 1966; Nicol et al., 1996a) then r vs. dplots seems to be very useful to determine the fractal dimension of a fault set. As a systematic difference between hanging wall subsidence and footwall uplift is often associated with the studied normal faults we set (r,) for the hanging wall rollover and (r2) for the wavelength of footwall uplift and constructed two different log-log plots.

The dmax vs. (r,) for 21 faults surveyed, are shown in Fig. 7b. The range of both dmax and (r,) is approximately three orders of magnitude, which is sufficient to define a linear relationship between these parameters. This relationship is expressed by S = 0.16 (6 is defined as d/r,) and by Eq. (1)

logdmax = l.O2*logr,-2.1. (1)

The c exponent of this equation is about 1, thus indicating that displacement and rollover length have increased proportionally during each growth increment of the faults, (s. Cowie


( UI) swamaa,s[ds!a


ALBANIA

Fig. X. The nwhern Hepirns t;~ul~ \et. Smplificd tectonic map (after Koukoclzas and Perwr. 1963. Perrier and

Koukouzas. 196X and ZOUI-o\. 1993) h, cros+wction. c Dlsplacemcnt-dist;lnce dugram of the i\rlsti fauh.

The western segment is composed of an array of smaller overlapping and underlapping segments along which displacement diminishes abruptly to zero.

If we take the rollover length (Y,) of the Konitsa fault to be 7 km (Fig, Xb). then by using the log-log plot of Fig. 7b (which relate r, and d for smaller faults) we estimate d to be equal to 900m. This displacement is nearly similar to the displacement of IOOOm derived by geological reconstruction and suggests that larger faults also remain scale invariant.


Furthermore, this estimate suggests that the segmented fault zone acts essentially as a single fault.

The Aristi fault in the southern part of the area have cross cutted a higher topography. Only a small part of the rollover structure is buried beneath young deposits which are mainly concentrated along a narrow depression in the central part of the fault. The fault zone itself consists of three, right-stepping en echelon fault segments. The displacement- distance plot (Fig. 8c), constructed using as marker horizon the contact between the Eocene limestones and the late Cenozoic flysch shows that maximum displacements occur at the center of the fault, near the overlap area between the two segments. The displacement line is slightly asymmetric showing the typical shape of interacting segments of faults described in several papers (Childs et al., 1995). If we consider the Aristi fault-zone as a single fault, and take r, = 5 km (Fig. 6), then a maximum displacement of 800 m along this fault can be predicted, by using the log-log plot of Fig. 7b.

As our fault data show a power law relationship, we can also estimate the total strain in the area (Marrett and Allmendinger, 199 1; Westaway, 1994). Based on the exponent C = 0.5 of the fault population in Fig. 4, the contribution to total extension from the non-observed fault fraction is 7%. Considering that the cumulative heave of the three faults in three examined sections is about 4 km we calculate a total extension of about 16%. Furthermore our fractal model predicts that considerable footwall uplift occurs (Fig. 7a) during extension. By the large displacements of more than 800m of the Konitsa and Aristi faults, uplift amounts in the order of 400 m are to be expected (Fig. 7~). Thus normal faulting essentially contributed to the high topographic relief or the area (2.438 m Gamila mountain crest, Fig.

8).

The Aliakmon fault zone

The Aliakmon fault zone extends for up to 70 km along the base of the Flambouro- Kamvounia mountain ranges and separates the greatly extended area of NW Macedonia from the more stable area of Thessaly (Fig. 1 and Fig. 9a). It is marked by large well preserved scarps associated with lacustrine, colluvial and alluvial deposits of middle to late Pleistocene age as well as with Holocene deposits. Geologic evidence and earthquake ruptures show a segmentation pattern and provide an excellent opportunity to establish kinematics of each one segment and to propose a fault growth model.

The Aliakmon fault zone comprises three en echelon fault segments. Two of them (the Sarakina and Servia faults) are linked to form a nearly straight fault zone whereas the Deskati fault merge with them to form a horsetail structure (Fig. 9a). We describe first kinematics of each one fault separately, and then propose a fault growth model.

(1) The Sarakina fault has been activated during the 13.5 1995 M = 6.6 earthquake to produce a 25 long NE trending surface rupture (details of the geometric pattern of this zone can be found in Mountrakis et al., 1995, Pavlides et al., 1995). Although this fault has been ruptured as a independent unit during the earthquake (‘earthquake segment’ sensu de Polo et al., 1991), it can be subdivided into three second order segments: A central segment (the Paleochori segment) is overlapped to the south by the Nisi segment and underlapped to the north by the Rymnio segment. Coseismic slip along the surface rupture has diminished westward from 20cm in the Rymnio trough 1Ocm in the Paleochori to zero in the Nisi segment. Slip on the Rymnio segment is confined to a well defined range front, whereas slip in the Paleochori and Nisi segments is transferred into a 3 km wide zone of diffuse fractures.

208 1‘. Doutsos and I. Koukouvelas


Coseismic ruptures of up to 2 km long, along these two segments, show vertical throws up to 10 cm and are antithetic to the main earthquake rupture. Antithetic small faults near to a tip point of larger faults is a common feature (Nicol et al., 1995) and may constitute a ‘break down zone’ of the Rymnion segment caused by inelastic deformation (S. Cowie and Scholz, 1992b). The maximum coseismic slip of 0.5m occurred at depth of 1Okm (Papazachos et al., 1995) along a principal nodal plane dipping 45”.

Commutative deformation preserved in the geologic structure is very similar to the coseismic deformation. Thus existing maps (Mavridis and Kelepertzis, 1974) and geophysical data (Iconomou, 1970) suggest that basin depths and rotation of the faulted strata vary along the strike of the Sarakina fault from east to west. The Rymnio segment to the east is 10 km long separating the Kamvounia range from the western part of the Aliakmon basin. Hanging wall subsidence along this fault is manifested by a PlioceneeQuaternary sedimentary wedge whose thickness is in excess of 600 m (Fig. SC). Height of fault escarpment across the 45” dipping fault surface is 900 m giving a displacement of 1500 m. A nearly similar value of 1600m can be estimated on the cross section of Fig. 9c (r, = 8 km and by using the log-log plot of Fig. 7b. This independent estimate of displacement with different methods suggests that displacement in the Rymnio fault is scale invariant.

In the Paleochori segment, further west, reduced hanging wall subsidence lead to the formation of shallow Quaternary sub-basins, which are filled with up to 200 m thick alluvial fan deposits (Fig. 9b). Often they are bounded by small antithetic faults whose strike changes from NE-SW to E-W (Fig. 9a). Height of fault escarpment across this segment is low and the estimated displacement of 300m differs greatly from the predicted value of 600 m (set r, =4 km in the Fig. 7b). This means that displacement along this fault segment is scale variant.

The Nisi segment, to the west, is 5 km long and overlaps the Paleohori fault. An along strike valley floor in the hanging wall of this fault tilts the earth’s surface slightly to the south and thin Quaternary sediments are accumulated in a small depression along with the low range front. The estimated displacement of 50m is not predictable by means of the log-log diagram of Fig. 7b and therefore is not fractal in nature.

From the above information, it is obvious that the Sarakina fault acts as a single segment. From east to the west each fault segment becomes younger, and shorter and shows smaller displacements. Similar fragmentation patterns characterize the active normal faults of the Dixie Valley in the basin and range province of western USA (Zhang et al., 1991). The d- r, ratio in the Sarakina fault is fractal only if we consider that r, of the Rymnio segment is representative for the whole fault length. Nevertheless the maximum displacement is not located at the fault’s central potion, as it is usually located on single faults, but occurs closer to the overlap zone with the Servia fault (s. also Fig. 10e). This is a reliable indicator of fault interaction between the Rymnio segment and the Servia fault, with the Servia fault not activated during the last earthquake. As the maximum displacement of 1700 m across the Sarakina fault has taken place in the past 6 Ma (Lower Pliocene) a long term slip rate of 0.3mm a-’ can be estimated. About 3500 seismic events with coseismic maximum displacement of 0.5m are required to accumulate this displacement. Thus the recurrence interval is in the order of 2000 years.

For the three-dimensional representation of the Sarakina fault, more geophysical data is required. It is very likely that the fault system is segmented in a vertical section, and therefore the continuous trace of the down dip width of faults in the cross sections of Fig. 9b and c represents only an oversimplification.


t’. */* $++

b rf= k0msta relay ramp lB= Vourinos barrier &..y

/’

e Fig. 10. Cartoon maps (at(d) to show growth model and (e) displacement-distance diagram of the .&Ii&non

fault zone.

However the high uplift/subsidence ratios described in our flexural model enable us to

assume, that at least a large part of the fault system have a planar geometry. Taking into account that most mesoscopic faults (Fig. 3), including the surface rupture of 13.5.95 earthquake dip about 6&80”, and that the large seismogenic faults in the northern Hepirus fault set and the Aliakmon fault system dips at 45-50”, we infer a shallowing of fault dip downward. Furthermore, the intentification of low angle to horizontal slip mechanisms of small aftershocks associated with the 1995 main event conforms to a listric geometry of the fault termination at the brittle-ductile transition in the crust (S. Makropoulos et al., 1996). In addition, the dense pattern of antithetic faults described in the area is considered by Bruhn and Schultz (1996) to be a reliable indicator of listric geometry. We propose therefore a three dimensional fault geometry characterized by a large central planar part and two small listric extremities.


(2) The Servia fault comprise two underlapping fault segments: The Velvendos segment to the west separates the higher crest of the Flambouro range (2.161 m) from that part of the Aliakmon basin which attain its maximum width (12km) (Fig. 9a). Although displacement along the Velvendos fault can not be recorded by geological observations it can be predicted if we set in the diagram of Fig. 7b the maximum width for the basin which is nearly equal to the rollover length. About 2100m of maximum displacement should have taken place across this fault. Furthermore a d-x profile was constructed to show the displacement gradient along this segment (Fig. 10). The asymmetric shape of this profile indicate interaction with the neighboring Rymnio fault (Fig. 10e). The Polyphyto segment to the east of the Sarakina fault forms lesser range fronts along which the Aliakmnon river flows to the Thessaloniki basin. As the hangingwall of this fault show a relatively high topography, affected only by erosion, we infer that lower displacements occurred across this segment.

(3) The Descati fault is 30 km long and consist of four fault segments along which structural setting varies from west to east considerably. The Katakali segment to the west separates a 5 km wide lacustrine basin of Pliocene age (Karpero basin) from older molassic sediments (Fig. 9a). Although hangingwall and footwall topography in these soft rocks is generally low, a 200 m high fault scarp can be observed. To the west this segment terminates with a small relay ramp associated with Pleistocene conglomerates and sandstones (Isomata relay ramp Fig. 9a). Unfortunately there aren’t geologic data concerning the depth of the basin and therefore only assumptions can be made. If we assume that the relationship between Y, and Max in the Katakali segment is a fractal one, then we extrapolate from Fig. 7b the dmax to be about 1000 m. The Vounasa segment in the east enters a wide area of crystalline rocks and shows a different hanging wall subsidence along its strike. Thus the western part of the hanging wall is buried below Plio-Pleistocene rocks whereas in the eastern part no deposition took place. If we take from the map Y, = 2 km, the maximum displacement along this segment may not excess 400m. The Tranovalton segment further east is associated with a 7 km long and 1.5 km wide depression filled with a very thin Pleistocene sequence. Height of fault escarpment across this fault is about 200m and therefore dmax is about 250m. The Mikrovalto fault overlaps the Tranovalto fault and connects the Deskati with the Sarakina faults. A 3 km long and 1.5 km wide zone with small antithetic faults is here formed in the hanging wall of N-S trending master fault with dmax about 150 m.

At the first stages of the deformation the Deskati fault grows independently from the Servia-Sarakina faults, as the Katakali segment initiated in the lower Pliocene, 20 km far from them (Fig. 10a). Later in the Upper PlioceneeLower Pleistocene the fault growth is propagated eastward by underlapping to form the Vounasa segment and give way to the Aliakmon river (Fig. lob). By further continuation of the deformation in the Upper Pleistocene the Tranovalton barrier were crosscutted by two shorter fault segments showing smaller displacements (Fig. lOc, d). As they merged with the Sarakina fault to the north, changed gradually their strike from NE- to N-S, suggesting interaction relationships.

The structural pattern of small underlapped fault segments is recently considered to be an indicator for rapidly propagated faults (Cartwright et al., 1996). This is the case only at the first growth stages of the Deskati fault, where it is independently nucleated from the Sarakina-Servia faults. Later westward propagation was inhibited to form a relay ramp and a wide interaction fault zone was formed further east as the fault was connected with the Sarakina segment.


(4) The,f&dr gro~~1I1 model. The results of our displacement analysis in the Aliakmon fault zone combined with the stratigraphic record of the sediments associated with this fault zone provide useful constraints on fault segmentation and allow us to present a fault

growth model in four stages (Fig. IO).

(a) Faulting initiated during the lower Pliocene independently at two locations along the Velvendos and Katakali segments of the fault zone (Fig. 10a). That are the locations where at present time the fault zone displays two separate displacement maxima (see

also Peacock and Sanderson, 1993). (b) By further continuation of the deformation in the Upper Pliocene time the Velvendos

segments continues to grow by out-of-plane bifurcation (Huggins ft u/.. 1995) to

form the Rymnio segment of the Sarakina fault. In the same time further south the Katakali subsegment propagated eastwards to form the Vounasa segment (Fig. lob).

(c) In lower Pleistocene times, the Rymnio segment were underlapped westwards by the

Paleochori and the Velvendos segments lengthened by eastward fault propagation. Further south the Katakali segment stopped to grow westwards by a small relay

ramp whereas eastwards the Vounasa segment was underlapped by the Tranovalton segment (Fig. 10~).

(d) At present time the Servia Sarakina fault Lone propagates at both sites with the formation of the Nisi and Polyphyto faults respectively, whereas the central part of

this zone was connected and interacted with the Deskati fault (Fig. 10d).

Furthermore, if we accept that steeply displacement gradients on fault segments are often the result of reduced tip line propagation (Nicol CI a/., 1996b) then we can recognize the

presence of the Agiophyllo tip which have restricted the Deskati fault from the undeformed molassic rocks to the west and the Vourinos barrier at the center of the Aliakmon fault

zone. The last consist of the Vourinos ophiolitic complex which prohibited the Servia fault to propagate westward. West of this barrier the Paleochori and Nisi segments enter in the late stage of deformation into an area of soft molassic sediments and caused the formation

of a wide zone of small antithetic faults. The abundance of segments with underlapped geometry in that area can be considered with (Cartwright et al., 1996) to be an indicator for rapid propagation.

We conclude that the first stages of fault growth are characterized by the nucleation and out-of plane bifurcation of two independently fault zones the Servia-Sarakina fault zones

to the north and the Deskati fault zone to the south, which grow laterally towards one

another. Only at later stages of deformation both segments were linked to form the Aliakmon fault zone (Fig. I Oe).

CONCLUSIONS

(1) Displacement analysis of small faults with dmax < I m reveals that faults with different size are scale variant. Data on d/h log log plots show a significant amount of scatter, e.g. faults of similar down dip width can have displacements that differ by a factor of 2.

(2) The cumulative frequency plot of displacements as well as the r vs. &ax plots of fault with lm < &ax < IOOm suggests that these faults are scale invariant. Large seismogenic faults with dmax = 1000 m such as the Petres, Konitsa and Aristi faults fit well in the curves of these diagrams indicating a scale invariant relationship.


(3) Kinematic parameters on reverse drag profiles are power-law. Their power law exponents are about 1 and indicate that displacement and rollover length, footwall uplift and wavelength of footwall uplift increased proportionally during each growth increment of the faults.

(4) The footwall uplift/hanging wall subsidence ratio is about l/2. (5) d-l profiles of large seismogenic faults are rarely symmetric in shape. At several sites,

these profiles are strongly influenced by segment linkage whereas in other cases they terminate abruptly on barriers showing steeply dipping displacement gradients.

(6) A growth model of the Aliakmon fault zone comprise four stages. The first two stages are characterized by the nucleation, tip line lengthening and out of plane bifurcation of the Servia-Sarakina faults to the north and the formation of the Deskati fault to the south. In the last two stages these faults continue to grow laterally towards one another until they meet at present time to form the Aliakmon fault zone.

(7) The Sarakina fault of the Aliakmon fault zone hosted the 13.5.1995 earthquake to produce a 25 km long surface rupture. Coseismic slip and cumulative displacement along this fault decreases from E to W. This decrease is discontinuous taking place along segments which become shorter and younger toward the same direction. The maximum displacement of 1700 m across this fault has been accumulated by 3500 seismic events, having a recurrence interval on the order of 2000 years.

Acknowledgements-Research is partly supported by the Public Power Corporation. Editing assistance of S. Pavlides and G. Cello comments and corrections improve the manuscript.

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Fractal analysis of normal faults in northwestern Aegean area, Greece

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