Interpretation of Low Seismicity in the Eastern Anatolian Collisional Zone Using Geophysical (Seismicity and Aeromagnetic) and Geological Data

Interpretation of Low Seismicity in the Eastern Anatolian Collisional Zone

Using Geophysical (Seismicity and Aeromagnetic) and Geological Data

M. NURI DOLMAZ,1 OMER ELITOK,2 and U. YALCIN KALYONCUOGLU1

Abstract—The eastern Anatolia is a continental collisional zone between the Arabia-Eurasia plates and is

currently being accompanied by the westward escape of the Anatolian crustal block to the west-southwest along

two major strike-slip fault zones, the NATF and the EATF. Although these major fault zones have experienced

historical earthquakes with moderate to high magnitude, notably there is a low seismicity zone on the active

EATF between the Bitlis and Poturge massifs. The low seismicity zone is characterized by thinner crustal

structure relative to its environs, the shallow Curie Point Depth (SCPD, ca 12–14 km in between 39–40�E and

38.5–39�N in the easternmost part of the Anatolian plate) and moderate to high b values (more than 0.7). We

consider that the shallow CPD and moderate to high b values in the low seismicity zone characterized by thinner

crustal area are closely related with the higher thermal structure of the crust, which most probably resulted from

crust-hot asthenospheric mantle interactions.

Key words: Eastern Anatolia, seismicity, Bitlis-Poturge massif, CPD, b value.

1. Introduction

The eastern Anatolia is a continental collisional zone that is currently being squeezed

and shortened between the Arabian and Eurasian plates. This collisional and contractional

zone is bounded by Pontide Belt (PB) in the N and and Bitlis-Poturge Suture Zone

(BPSZ) in the S (Fig. 1). The contraction and thickening of the crust to ca 50–52 km

(DEWEY et al., 1986; PEARCE et al., 1990) in the collisional zone has been accompanied by

tectonic escape of most of the Anatolian crustal block (Anatolian plate) to the west-

southwest towards the Aegean-Cyprean arc system by major strike-slip faulting on the

right-lateral North Anatolian Transform Fault (NATF) and left-lateral East Anatolian

Transform Fault (EATF) (SENGoR and YILMAZ, 1981 and references therein; SENGoR et al.,

1985; DILEK and MOORES, 1990; HUBERT-FERRARI et al., 2003). Therefore, widespread

seismic activity in eastern Anatolia is related with the ongoing collision and crustal

escape tectonics. Although seismic activities with moderate to high magnitude are

confined along the major fault zones (EATF, NATF, BPTZ), there is notably a low seismic

1 Department of Geophysical Engineering, Suleyman Demirel University, 32260 Isparta, Turkey. E-mail:

[email protected] Department of Geological Engineering, Suleyman Demirel University, 32260 Isparta, Turkey.

Pure appl. geophys. 165 (2008) 311–330 � Birkhauser Verlag, Basel, 2008

0033–4553/08/020311–20

DOI 10.1007/s00024-008-0307-yPure and Applied Geophysics

zone from the easternmost part of the Anatolian plate to the Arabian foreland through a

gateway called metamorphic gap between the Bitlis and Poturge massifs. This area is also

characterized by thinner crustal structure relative to its environs. In light of geological

and geophysical data, this study aims to investigate the low seismicity of the region

although occurring in a tectonically active region. For this purpose: i) The CPD map of

the study area was prepared from spectral analysis of aeromagnetic data, ii) the crustal

thickness map of the same area was redrawn on the basis of estimations of GoK et al.

(2007), which is based on receiver functions from recordings of 29 broadband stations,

and iii) the b parameters were evaluated and hence the b-value map of the study area was

prepared from the earthquake data.

2. Regional Tectonic Setting

The eastern Anatolian collisional zone, formed by convergence and collision of the

Arabia-Eurasia plates, takes place in the Alpine-Himalayan orogenic belt. This collision

zone, which consists mainly of a collage of fragments of oceanic and continental crusts, is

bounded by the Arabian foreland by a suture zone (Bitlis-Poturge suture) in the south

(Fig. 2). The collage of fragments of oceanic and continental crusts forms a nappe

succession running along the Bitlis-Poturge suture zone. YILMAZ (1993) and YIGITBAS

et al. (1993) separated the nappe successions into two main zones: i) The nappe zone

consisting of two nappe stacks (the upper nappe represented mainly by Bitlis and Poturge

Figure 1

Simplified tectonic map of Turkey and surrounding area (modified from BOZKURT, 2001), EACP: Eastern

Anatolian Contractional Province, CAOP: Central Anatolian Ova Province, WAEP: Western Anatolian

Extensional Province, NATF: North Anatolian Transform Fault; EATF: East Anatolian Transform Fault;

NEAFZ: North Eastern Anatolian Fault Zone; DSFZ: Dead Sea Fault Zone; K: Karlıova.

312 M. Nuri Dolmaz et al. Pure appl. geophys.,

metamorphic massifs; the lower nappe is characterized by the slices of polyphase

metamorphic ophiolitic assemblage and the Maden Group), ii) the imbricated zone

sandwiched between the Arabian platform and the nappe zone. The nappe successions,

including both the nappe zone and the imbrication zone of YILMAZ (1993), form together

the Bitlis-Poturge Thrust Zone ‘‘BPTZ’’ (Fig. 2). Along the BPTZ, all the allochthonous

units overlie tectonically the autochthonous Arabian foreland units. Compositionally

Figure 2

Simplified geological map of the eastern Anatolian region (modified from BINGOL, 1989).

Vol. 165, 2008 Low seismicity in Eastern Anatolian 313

variable and young volcanic units that erupted during the Neogene and Quaternary times

cover most of the eastern Anatolian region and the Arabian foreland (PEARCE et al., 1990;

YILMAZ et al., 1998; KESKIN, 2003; KESKIN et al., 2006). The Bitlis massif consists of a

metamorphic core intruded by Precambrian granites and a cover sedimentary pile of

Paleozoic-Mesozoic rocks (CAGLAYAN et al., 1984). On the other hand, GoNCuOgLU and

TURHAN (1984) interpreted that the Bitlis metamorphic belt consists of numerous tectonic

slices thrust over one another and ophiolite obduction-related deformation and

compression developed during Upper Cretaceous in the metamorphic belt. CAGLAYAN

et al. (1984) observed that post-metamorphic diabase dykes associated with upper

Cretaceous-middle Eocene volcanism cut across the crystalline basement. CAGLAYAN

et al. (1984) and GoNCUOgLU and TURHAN (1984) drew attention to the similarities of

Paleozoic and Mesozoic rocks of the Bitlis metamorphites with the units of correspond-

ing ages in the Arabian autochthon. They also stated that the Malatya metamorphics

consist mainly of metamorphosed platform carbonates. The Keban metamorphics in the

north of the Malatya and Poturge massifs are composed mainly of the metamorphosed

platform carbonates with amphibolites, sandy limestones, calc-schists (YAZGAN, 1984)

and are cut by upper Cretaceous latite, trachylatite dykes (BINGOL, 1984).

The Arabian foreland is composed of a continuous stratigraphic sequence of mainly

shelf sediments of early Paleozoic to Miocene age resting on a Precambrian basement

(PEARCE et al., 1990; HALL, 1976; YILMAZ, 1993; YIGITBAs et al., 1993). KARACADAG

volcano in the Arabian foreland erupted since the Pliocene along a N-S trending set of

fissures and craters, spatially associated with the nearby Akcakale graben (Fig. 2)

(SENGoR et al., 1985 and references therein; PEARCE et al., 1990). PEARCE et al. (1990)

interpreted that small degrees of stretching might have caused melting of metasomatized

lithosphere by perturbation of the geotherm by heat from the upwelling of hot

asthenosphere. In relation with this, they stated that Akcakale graben, a small rift

structure with N-S normal faults and located to the SW of Karacadag, is an indicator of

E-W extension, although this rift appears to act, at least in part, as a transfer structure

between some of the outer thrusts of the Arabian foreland sedimentary sequence.

3. Methodology of Magnetic Basement Depth Estimate

The depth of magnetic sources can be estimated by spectral analysis. If the basement

rocks are magnetized, the base of magnetic sources (zb) is assumed to be the Curie Point

Depth (CPD). At temperatures greater than the Curie point, the dominant magnetic

mineral (ca. 580�C for magnetite) in the crust passes from a ferromagnetic state to a

paramagnetic state (NAGATA, 1961). Beneath the CPD the lithosphere shows virtually

nonmagnetic properties.

The Curie Point Depth (CPD) has been estimated by spectral analysis of

aeromagnetic data to understand the thermal structure of the crust and distribution of

magnetic basement depths in various tectonic settings around the world (e.g., VACQUIER


and AFFLECK, 1941; SMITH et al., 1974, 1977; BHATTACHARYYA and LEU, 1975; BYERLY and

STOLT, 1977; SHUEY et al., 1977; BLAKELY and HASSANZADEH, 1981; CONNARD et al., 1983;

OKUBO et al., 1985, 1989; BLAKELY, 1988; OKUBO and MATSUNAGA, 1994; HISARLI, 1995;

TSOKAS et al., 1998; TANAKA et al., 1999; BADALYAN, 2000; STAMPOLIDIS and TSOKAS,

2002; DOLMAZ et al., 2005a, b; ATEs et al., 2005; AYDIN et al., 2005; LıN et al., 2005;

BEKTAs et al., 2007). In general, the CPD is classified as shallow CPD (< ca. 16–18 km)

and deep CPD (> ca. 16–18 km). The CPD is in close relationship with the crustal

thickness, lithospheric scale geologic structures, depths of brittle and ductile deformation

zones, crust-mantle interactions and magmatic events.

The method used to examine the spectral knowledge included in subregions of

magnetic data was developed by OKUBO et al. (1985) and TANAKA et al. (1999), and

resembles the method of SPECTOR and GRANT (1970). SPECTOR and GRANT (1970) showed

that the expectation value of the spectrum of an ensemble model was the same as that of a

single prism with the average parameters for the collection.

CPD (zb) is briefly estimated in two steps as suggested by BHATTACHARYYA and LEU

(1975), OKUBO et al. (1985), and TANAKA et al. (1999). The first is the depth to the

centroid (z0) of the deepest magnetic source from the slope of the longest wavelength part

of the spectrum divided by the radial frequency,

lnP sj jð Þ1=2

sj j

" #¼ ln A� 2p sj jz0; ð1Þ

where P sj jð Þ is the power density spectrum of the anomaly, sj j is the wavenumber, and A

is a constant. The second step is the estimation of the depth to the top boundary (zt) of that

distribution from the slope of the second longest wavelength spectral segment,

ln P sj jð Þ1=2h i

¼ ln B� 2p sj jzt; ð2Þ

where B is a constant.

The CPD estimates have been carried out based on the following three stages: 1)

Dividing into overlapping square subregions, 2) calculating of the radially averaged log

power spectrum for each subregion, 3) estimating of the CPD (zb) from the centroid (z0)

and the top depth (zt) estimated from the magnetic source for each subregion using the

following equation;

zb ¼ 2z0 � zt: ð3Þ

At calculations, the horizontal dimension of a magnetic source must be considerably

much larger than the top depth (zt), and the radial averages of magnetization and

geomagnetic field direction must be constant (LıN et al., 2005). Furthermore, because the

shape of the spectrum curve is independent of the inclination and declination of the local

geomagnetic field and the magnetization, the spectrum slope would not be changed even

if the distribution of magnetization is not random in any area (GARCIA-ABDESLEM and

NESS, 1994).


4. Aeromagnetic Data Processing

The aeromagnetic data of the study region were obtained from the General

Directorate of Mineral Research and Exploration (MTA) of Turkey, collected along flight

lines spaced at 1–3 km profile intervals at an elevation of 600 m above ground level

spanning a in period of 1978–1989. We subsequently carried out the International

Geomagnetic Reference Field (IGRF) correction on the data utilizing the program of

MALIN and BARRACLOUGH (1981) for the year 1982.5. The total field aeromagnetic

anomaly data were produced after the removal of the IGRF and the data were interpolated

to a regular grid of points with 2.5-km spacing. The total magnetic field data were then

transformed into the (north) magnetic pole (BARANOV, 1957), utilizing the FFTFILL

program (HILDENBRAND, 1983).

The CPD estimations (zb) require the deepest magnetic sources and are obtained from

using wavelengths longer than 10 km (TANAKA et al., 1999; STAMPOLIDIS and TSOKAS,

2002; DOLMAZ et al., 2005a, b). In order to emphasize the effect of deep sources, the small

wavelength anomalies must be removed from the anomaly data. For this purpose, a

simple band-pass filter (full pass 10–65 km) was designed from the response function of

the power spectrum of the reduced pole data and applied to these data by using the

FFTFIL (HILDENBRAND, 1983) again (Fig. 3). The filtered data used for the CPD

estimation are illustrated in Figure 4.

We have tried to estimate the CPD of the study area from the band-pass filtered data

employing the methodology above, using divided blocks of size 90 9 90 km2 and we

overlapped fifty percent with the adjacent block because of the limited depth extent of

crustal magnetization. MAUS et al. (1997) already implied that areas of the magnetic data

Figure 3

Azimuthally averaged log power spectrum of the reduced pole total-field aeromagnetic data. The box shows the

band-pass filter amplitude response. The response function is also shown.


must not exceed 100 9 100 km. We first calculated the radially averaged log power

spectrum of each block utilizing once again the FFTFIL (HILDENBRAND, 1983). Examples

of power spectra of magnetic anomaly data are shown in Figure 5, where z0 and zt are

obtained by computing power spectrum slopes of the longest wavelength part of the

spectrum and the second longest wavelength part of the spectrum, respectively. From the

z0 and zt, the CPD (zb) is calculated for a divided block by means of the Equation (3).

The CPD contours constructed from the CPD estimates by using the standard gridding

routine have been shown on the crustal thickness map of the study area (Fig. 6).

5. Crustal Structure and Seismicity

Based on the Eastern Turkey Seismic Experiment (ETSE) project (SANDVOL et al.,

2003), crustal thicknesses of E Turkey have been estimated by using receiver functions

from recordings of 29 broadband stations (ZOR et al., 2003 and GoK et al., 2007). Using

the moho thickness estimates from Figure 7 in GoK et al. (2007) in the region, we gridded

Figure 4

The band-pass filtered map, used for the CPD calculations. Contours every 50 nT. Locations of the major

structures are also shown. NATF, North Anatolian Transform Fault; EATF, East Anatolian Transform Fault;

BPTZ, Bitlis Poturge Thrust Zone; Mal, Malatya; Ady, Adıyaman; Elz, Elazıg; Tun, Tunceli; Erz, Erzincan;

Bin, Bingol; Diy, Diyarbakır; Mar, Mardin; Bat, Batman.


and extrapolated them to obtain the crustal thickness map of the study area (Fig. 6). In the

map, the crustal thickness of the study area becomes thinner from north (ca. 42–44 km) to

south (ca. 36–38 km) relative to its eastern (EAHP) and western sides (central Anatolian

region), extending into the Arabian foreland through the metamorphic gap between the

Bitlis and Poturge massifs (Fig. 6). Seismic data have also shown that crustal thicknesses

in the north of Bitlis and Poturge metamorphic massifs reach ca. 46 km.

Figure 5

Examples of power spectra defined from the band-pass filtered anomaly data to estimate zt and z0 at 39.15�N;

39.85�E. Dots correspond to values of power spectrum. a) The top of the magnetic basement zt = 3.19 km is

obtained by fitting a straight line through the high wave number portion of the data. b) The depth of the centroid

z0 = 9.8 km is obtained by fitting a straight line through the low wave number portion of the data.


Bougouer gravity anomalies in E Turkey are as low as -160 mgal (ATEs et al., 1999).

BARAZANGI et al. (2006) recalculated the gravity anomalies using the Moho structure from

ZOR et al. (2003). Their residual gravity anomaly map outlined by subtracting the

observed gravity values from their calculated ones indicates that low gravity residuals in

south of the Bitlis-Poturge suture result from the lower density material at sub-Moho

depths. They interpreted that crustal density variations are the main source of the low

gravity anomalies. Since very low Pn velocities and very high Sn attenuation exist in the

uppermost mantle, the low gravity residuals are caused by the presence of asthenospheric

material at sub-Moho depths (BARAZANGI et al., 2006).

The earthquakes that occurred during 1964 to 2006 were plotted on the CPD variation

map with the aim of comparing the seismicity and the thermal structure of the study area

(Fig. 7). The earthquakes in the study area have been investigated by several studies (e.g.,

ERCAN, 1982; EYIDOgAN, 1983; OSMANsAHıN et al., 1986; PERINcEK et al., 1987; TAYMAZ

et al., 1991; PINAR, 1995; CETIN et al., 2003; SANDVOL et al., 2003; TURKELLI et al., 2003;

ZOR et al., 2003; GURBUZ et al., 2004; OVER et al., 2004; HORASAN and BOZTEPE-GuNEY,

2007). The seismic data were taken from the ISC (International Seismological Center)

catalogues and the earthquakes of magnitude C 3 were used only. Overall, the seismic

activity is confined along fault zones (EATF and NATF) in the study area (Fig. 7)

however a scatter is observed over a wide region on the northwestern side of the study

area. Dense earthquakes between the latitudes of 38 �N and 38.8�N are thought to be

associated with the Bitlis and Poturge massifs which coincide with the continental

Figure 6

Map showing the Moho depth variation of the study area in grey scale as obtained by GoK et al. (2007). The

CPD contours in thick lines are also plotted on the crustal map. Contours every 2 km. Names of the features and

places as in Figure 4.


collision of the Arabian-E Anatolian Plates (Fig. 7). Moreover, there is hardly any

seismic activity approximately south of the latitude 38�N.

The E-W trending portion of the shallow CPD (SCPD) region in the east of Elazıg lies

in a low seismicity zone (LSZ), which is between the two high seismic activity zones

extending along the Poturge massif and the Bitlis massif (Fig. 7). The earthquakes on the

Poturge massif near the southwest of Elazıg City occurred at the near western edge of the

inferred thermal structure (SCPD; Fig. 7). Furthermore, the eastern edge of the inferred

thermal structure (SCPD) has been damaged repeatedly by large earthquakes (the

earthquakes around the City of Bingol, south and southwest of Bingol in Fig. 7).

Earthquakes also occurred around the eastern part of Malatya City close to the margin of

the inferred thermal structure (SCPD). It is interesting to note that the approximately E-W

trending part of the shallow CPD region (SCPD) cuts the EATF. There is hardly any

seismic activity in the overlap region (SCPD), while intense seismic activity has been

recorded on this fault zone (EATF) beyond the overlap area (Fig. 7).

Vertical distribution of the focal depths of the earthquakes, position of the CPD,

topography and Moho depth are shown in three cross sections taken along 50-km wide,

NE-SW, NW-SE trending profiles (see location of profiles in Fig. 7).

Figure 7

Seismicity of the study area for the period 1964–2006. The CPD contours are also plotted at 2 km intervals on

the map for comparison. Thickened lines indicate locations of the cross sections used in Figure 8. LSZ, low

seismicity zone. Names of the features and places as in Figure 4. See text for explanation.


Figure 8a: Profile involves the NE-SW directed cross section with 320 km between

latitudes 38�N–40�N. The cross section indicating the CPD clearly illustrates the shallow

Curie isotherm depth (SCPD) between longitudes 39.4�E and 40�E. This area is

characterized by a thinner crust of ca. 38–40 km, corresponding to the Low Seismic Zone

(LSZ) between the Poturge massif and Bingol surrounding area. The portions of the

profile between 0 to 80 kms and more than 145 km are characterized by deep CPD

estimates reaching a maximum depth of ca 19 km and 2.8 km topographic height. In

between 80 to 145 kms of the profile, the shallower CPD (SCPD) estimates of about

13 km are observed in the Low Seismic Zone (LSZ) between the Poturge massif and

Bingol towns, where intense earthquake distributions are observed (Fig. 8a).

Figure 8b: First part of profile indicates the NE-trending and second part of profile E-

trending cross sections with 250 km between longitudes 38�N–39�N. The cross section

clearly reveals the presence of one shallow maxima (SCPD) reaching a maximum depth

of ca. 13 km and two deep maxima; one in part nearly 0–50 km of the profile and the

other in part nearly 150–200 km of the profile. The first deep maximum is ca. 18-km deep

and is located in the south, beneath the Poturge massif. Its topographic range is from 1 to

2 km. The second one to the east is ca. 18-km deep, corresponding to the Bitlis massif.

The two negative spikes on the CPD are consistent with the suture structures (the Bitlis

and Poturge massifs) and intense earthquake distributions. In between longitudes of

39.3�E and 39.9�E on the profile, a shallow CPD (SCPD) zone reaching a maximum

shallow depth of ca. 13 km is located in the LSZ, forming at a height of 0.8–1.2 km

topography.

Figure 8c: Profile involves the NNW-SSE directed cross section with 200 km

between the longitude 40�E–41�E. The cross section generally shows the deep CPD

estimates, ranging depths of ca. 16 to 20 km. Distribution of the earthquakes shows

intense active seismicity under the wrench tectonics of E Anatolia. The portion of the

profile more than 80 km (latitude greater than 38.7�N) is characterized by more intense

earthquake activity than the Bitlis massif at 0–80 km part of the profile. The topography

in the north of the Bitlis massif varies between 1.3 and 2.4 km, but around Bitlis massif

between 0.7 and 1.8 km.

6. Analysis of b-Values

For the calculation of the seismicity and seismic hazard parameters, some criterions

are considered for a selection of the earthquakes from the catalogs. The magnitude of the

historical period of earthquakes usually suffers from large errors which cause many

problems in seismicity and seismic hazard evaluation. For this purpose, we selected only

instrumental data between 1964 and 2006, magnitudes C 3 from the ISC catalogues. The

instrumental data, which are magnitude and earthquake locations in this period, are well

determined than the historical period earthquakes since being available to be recorded

and processed by many seismological station or networks.


It is widely accepted that the b parameter in seismology is related to the tectonic and

earthquake physics: i) The b value which is closely related with rock type, state of stress,

the ductility of rock, decreases when shear stress increases (SCHOLZ, 1968; WYSS, 1973;

URBANıC et al., 1992; CHAN and CHANDLER, 2001; SCHORLEMMER et al., 2005), ii) the areas

Figure 8

The cross sections taken along the profiles in Figure 7, showing the variation of CPD, topography, Moho depth

(M) and the hypocenters of the earthquakes that occurred between 1964–2006. a) Profile 1 along EATF, b)

Profile 2, and c) Profile 3. SCPD, shallow CPD; LSZ, low sesimicity zone. Length of profile is in km and also in

degree sheet format. Note that thermal anomaly (SCPD) region is located between the two intense seismic zones

and is characterized by low seismic zone (LSZ). See Figure 7 for the location of the cross sections.


of greater material of heterogeneity are characterized by high b value but the areas of low

degree of material heterogeneity by low b value (MOGı, 1962; HATZıDıMıTRıOU et al.,

1985; MAIN et al., 1992; MANAKOU and TSAPANOS, 2000), iii) the low b value is related to

the spacing and clustering properties of epicenters or distribution of fault segments

(HUANG and TURCOTTE, 1988; ONCEL et al., 1996; LAPENNA et al., 1998; NANJO et al.,

1998), and iv) volcanic regions and areas of high thermal gradients are characterized by

high b-values (WARREN and LATHAM, 1970; WIEMER et al, 1998; KATSUMATA, 2006).

The choice of the statistical method used for deriving the a and b values in the

Gutenberg-Richter (G-R) relation is a critical consideration. The two empirical constants

are commonly derived using the linear regression method on grouped magnitude data. A

detailed analysis of the dependence of the b value was presented by BENDER (1983) on the

interval size, maximum magnitude, sample size, and the data fitting techniques. Because

of the different assumptions regarding these parameters, considerably different b values

and the associated standard errors may be obtained from the same data set (BENDER,

1983). While no single data fitting technique can yield accurate b values from an

inherently incomplete seismic catalogue, it is important to ensure that the spatial

variations in the b value obtained from different data fitting techniques are consistent with

each other (KIJKO, 1988; CHAN and CHANDLER, 2001).

The relative size distribution of earthquakes is an essential input parameter needed to

perform probabilistic hazard analysis. The basic well-known equation of Gutenberg-

Richter (G-R) relation (ISHIMOTO and IIDA, 1939; GUTENBERG and RICHTER, 1944), one of

the well-fitted empirical relations in seismology, presents the frequency of occurrence of

earthquakes as a function of magnitude:

log NðMÞ ¼ a� b �M ð4Þ

where N is the cumulative number of shocks within a magnitude interval M ± DM,

mathematical parameters a and b depend on the seismicity rate which varies greatly from

region to region and the properties of the focal material, respectively.

We used the Kaltek method (KALYONCUOGLU, 2007) to make a b value map of the study

area. Since the method was described in KALYONCUOGLU (2007), we only introduce the

method summarily in the context of this paper. KALYONCUOGLU (2007) suggests that some

types of deceptive illustrations on the seismicity map arose from the specific magnitude

range: i) Suddenly decreasing at the number of earthquakes in a magnitude interval from

minimum to maximum, ii) closest occurrence frequencies in a wide magnitude interval,

and iii) accumulation of the earthquakes around the minimum magnitudes. Due to this

handicap, KALYONCUOGLU (2007) constituted one assumption ‘‘the a value in the G-R

relation demonstrates exponential distribution of the earthquakes in zero magnitude’’ and

one hypothesis ‘‘the a value calculated from whole region data set can be accepted as a

constant value for the calculation of new b values belonging to the each subregion which

are included by the main region’’ for the calculation of new b values. On the other hand,

the number of earthquakes which have zero magnitude is equal to the constant value for

each subregion or every point of the whole region using the following equation;


b ¼ a �

PNd

i¼1

Mi �PNd

i¼1

logðNiÞ � Nd �PNd

i¼1

Mi � logðNiÞ� �PNd

i¼1

logðNiÞ �PNd

i¼1

M2i �

PNd

i¼1

Mi �PNd

i¼1

Mi � logðNiÞ� � ð5Þ

where Nd is the number of data.

In order to produce a b value map of the study area, the entire region was divided into

grids with a spacing of 0.1� in latitude and longitude. The minimum critical number of

earthquakes was accepted as 5 shocks in each circle which were drawn around each grid

point. For each grid point, a and b values were calculated by using the Kaltek method.

The grid point centers were illustrated with a square shaped in Figure 9. Therefore we

constituted the b value map of the study area from the calculated b values using a

standard gridding routine (Fig. 9). In the southern part of the study area, we could not

calculate the b values due to the unavailability of five shocks in each circle. Because in

general it is accepted that 5 to 10 events are sufficient to reflect the general tectonic

features of a region in the Kaltek method.

Firstly, we calculated a and b values for every subregion using the classic method.

These are amin = 1, amax = 6.3, aavr = 3, bmin = 0.18, bmax = 1.5, bavr = 0.71.

Standard deviation of b value is calculated as 0.37. We accepted the constant a value

as aavr = 3 that calculated in the classic manner. Then the minimum, maximum, and

Figure 9

Distribution of b values for earthquakes in the study area with M C3 which occurred from 1964 to 2006. In each

sample N C 5 events are selected at nodes separated by 0.1�90.1�. The CPD in thick lines are indicated also by

contours every 2 km on the b-value map for comparison. Names of features and places as in Figure 4. In the

southern part of the study area, the b values were not calculated due to unavailability of 5 shocks in each circle.


average b values were determined as bmin = 0.51, bmax = 0.96, and bavr = 0.7 according

to the Kaltek method. Standard deviation of b value was calculated as 0.11. It is observed

that the Kaltek method decreases standard deviation in b value. These results show that

the calculated average b value according to the Kaltek method and the average b value,

which is calculated using the classic way, are of the same value (bavr = 0.7). The Kaltek

method changed only minimum and maximum b values determined from the classic way

but did not change the average.

The b value map shows that the b values are not homogeneous in the study area. The

average b value was determined as 0.7 and shown in dashed lines in Figure 9. The b

values above and under this average value are classified as ‘‘high b value’’ and low b

value’’, respectively. To compare the b values and the CPD, we posted the CPD contours

in 2-km interval on the b value map of the study area (Fig. 9). Two low b value anomalies

are observed in areas covering the easternmost part of the Anatolian plate to the Bitlis

massif and the Poturge massif. On the other hand, the northern and southern sides of the

Poturge massif and some part of the Bitlis massif are characterized by high b value

anomalies. The area between the Bitlis and Poturge massifs, and also its northern

and southern sides in a narrow zone are characterized by moderate to high b values

(around 38.5�N; 39.5�E). The observed high b-value anomalies mostly result from

shallow (*0–10 km) and small magnitude (M B 4) earthquakes.

7. Discussion and Conclusions

The eastern Anatolian contractional zone (EACZ), consisting of a collage of

fragments of oceanic and continental crust, is a current active collisional convergent

zone that is still being squeezed between the Arabian and Eurasian plates. This

collisional zone is also being experienced as a compressional-extensional tectonic

regime conducted by the westward extrusion of the Anatolian plate along the right-

lateral North Anatolian Transform Fault (NATF) and left-lateral East Anatolian

Transform Fault (EATF) (SENGoR and YILMAZ, 1981 and references therein; SENGoR

et al., 1985; HUBERT-FERRARI et al., 2003). Although, seismic activities with moderate to

high magnitude are confined along the major fault zones (EATF, NATF, BPTZ),

notably there is a low seismicity zone (LSZ) between the Bitlis and Poturge massifs on

the EATF. The crust consisting mainly of accreted continental and oceanic lithospheric

materials in the LSZ is thinner than its eastern and western sides. This thinner crustal

area also displays a ductile or rigid-ductile behavior with respect to its environs, and is

characterized mainly by shallow CPD and moderate to high b values. Various

interpretations have been suggested for the b values. For example i) the mechanical

structures of materials and stress conditions (MOGı, 1967), ii) rock type and ductility of

rock (SCHOLZ, 1968), iii) spacing or clustering properties of epicenters or distribution of

fault segments (HUANG and TURCOTTE, 1988; LAPENNA et al., 1998; NANJO et al., 1998),

iv) volcanic regions and areas of high thermal gradients (WARREN and LATHAM, 1970;


WIEMER et al, 1998; KATSUMATA, 2006). Most of all these interpretations may be

considered for the eastern Anatolian region. The seismogenic layer of the area (the

LSZ) between the Poturge and Bitlis massifs and also its northern sides seem to be

rather shallow and therefore may not accumulate enough stress for a large earthquake to

occur. In other words, deformation takes place in a ductile manner. ITO (1999)

interpreted that the changes in the thickness of the seismogenic layer strongly depend

on temperature. HYNDMAN and WANG (1993) stated that the transition from seismic to

aseismic zones is related to a thermal effect. The seismic-aseismic boundary is also

thought to be related to the brittle-ductile boundary in the crust (KOBAYASHı, 1976;

SıBSON, 1982; DOSER and KANAMORI, 1986; ITO, 1990; TANADA, 1999). In the study area,

dense earthquakes seem to occur in areas where the lateral gradients of the CPD are

steep. These areas may correspond to the boundaries between high and low thermal

regions of the crust. Therefore, the variations in the seismic activities may be closely

related to thermal structures of the crust which is responsible for Curie isotherm

distribution (CPD). Consequently, we conclude that hot asthenospheric mantle-crust

interactions in the LSZ resulted in the shallow CPD, the high b value, the high thermal

structure and hence ductile behavior of the crust.

Acknowledgements

We thank the International Seismological Center (ISC) for use of the earthquake data

and MTA of Turkey for use of the aeromagnetic data. We would like to thank anonymous

reviewers for their thorough and constructive review of the manuscript.

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(Received May 7, 2007, accepted September 15, 2007)

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Interpretation of Low Seismicity in the Eastern Anatolian Collisional Zone Using Geophysical (Seismicity and Aeromagnetic) and Geological Data

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