attenuation of p and pp waves in vrancea area – romania
TRANSCRIPT
ORIGINAL ARTICLE
Attenuation of P and pP waves in Vrancea area – Romania
Marian Ivan
Received: 29 December 2005 /Accepted: 21 September 2006 / Published online: 19 December 2006# Springer Science + Business Media B.V. 2006
Abstract The seismic attenuation in the Vrancearegion (Romania) is investigated from teleseismicrecordings of P and pP waves during the four major,intermediate-depth Romanian events that occurredsince the onset of digital instrumentation. Moststations are located in Canada and in the UnitedStates, being equipped with a variety of sensors,especially short-period ones. The amplitude spectralratio method is used, assuming no frequency depen-dence of the QP factor in the range 0.2–2 Hz. Noapparent correlation between the derived attenuationvalue and the type of recording sensor is observed.Lateral variations of the attenuation are obtained, witha very low QP area (values down to 33) located in thenorthwestern part of the Vrancea seismogenic vol-ume. For the stations with different azimuth angles inrelation to the epicentral area, QP values routinelyexceed 200. Most likely, the low attenuation valuesare related to an upwelling mantle material locatedimmediately beneath the crust, but limited in depth toat least 100 km.
Key words attenuation . P and pP waves . spectralratio .QP factor . Vrancea (Romania)
Introduction
Located in the Southeastern Carpathians (Figures 1and 2), the intra-continental seismic area of Vrancea(Romania) is mainly characterised by hypocentershighly confined to a quasi-prismatic, near-verticalvolume in the depth range of 85–220 km. Crustalseismicity is negligible and a seismic gap is presentaround the 40- to 60-km depth. Both the physicalmechanism of the intermediate-depth events and thenature of the seismogenic volume are still a questionof debate. Various geodynamic models have beenproposed to explain Vrancea seismicity, assuming apaleosubduction process of an oceanic slab (Fuchs et al.1979), basalt–eclogite phase changes (Oncescu 1980;Enescu 1985), mantle delamination (Gîrbacea andFrisch 1998; Chalot-Prat and Gîrbacea 2000), roll-back, the detachment and/or the final stage of platebreak-off of a subducted continental lithosphere(Csontos 1995; Mason et al. 1998; Seghedi et al.1998; Linzer et al. 1998), eventually followed by roll-back and steep dipping in the upper mantle (Gvirtzman2002). An active continental subduction beneath theCarpathian arc (Enescu and Enescu 1993) or thesubduction of a normal oceanic lithosphere, detachedfrom the Eastern European Plate, possibly caused by abreaking-off process from the Moesian Platform
J Seismol (2007) 11:73–85DOI 10.1007/s10950-006-9038-7
M. Ivan (*)Department of Geophysics,University of Bucharest, 6 Traian Vuia str.,020956 Bucharest o.p.37, Romaniae-mail: [email protected]
M. IvanSchool of Ocean and Earth Sciences,Tongji University, 1239 Siping Rd.,200092 Shanghai, China
(Wortel and Spakman 2000), have also been invoked.An actual tearing-off of a formerly subducted fragmenthas also been suggested (Sperner et al. 2001), while arefined gravitationally sinking stretched body modelhas been discussed (Cloething et al. 2004), partially inrelation to an evolutive process from an activesubduction to a passive sinking due to gravity, aspredicted by stress modelling (Ismail-Zadeh 2003).
Tomographic studies (Oncescu 1984; Fan et al.1998) indicated a higher velocity, near-vertical anom-aly, approximately located between 80 and 250 kmdepth, with a small low-velocity body placed in thenorthwestern part (Martin et al. 2005). The maximumvelocity perturbation is about 4%.
Macroseismic studies regarding some large events inVrancea outlined a clear ‘kidney’ shape of the iso-seismal lines, indicating major differences on both sidesof the Carpathian Arc (Enescu 2003). A strongasymmetric pattern, with amplitudes decreased by afactor of 20 along the inner volcanic chain withrespect to the foreland platform has been obtained bythe analysis of seismograms (Popa et al. 2005) andexplained by an upper mantle attenuation volumeplaced in the NW portion of the Vrancea seismic body.
However, the presence of a near-surface Neogene–Quaternary magmatic material or the existence about100 km deep of a body slowing down the seismicwaves have been completely discarded by Enescu(2003), who considers that the focal mechanism is themain factor responsible for the pattern of the iso-seismal lines. Using recordings from the University ofUtah and Washington University networks performedduring the major Romanian earthquakes of 30 August1986 and 30 May 1990, QP values down to 50 havebeen obtained in a small area located in the NW ofVrancea (Ivan 2003a), possibly indicating a spatialvariation of attenuation. Other results have beenobtained (Sudhaus and Ritter 2005) from four tele-seismic phases of two events underneath Japan andNew Guinea recorded from a seismic refractionexperiment along a N–S profile across the Vrancearegion, most likely correlated to three-dimensionalattenuating structures at deep lithospheric levels.Using a mathematical similarity between the attenua-tion parameter and the propagation time of the P-wave, the VELEST code (Kissling et al. 1994) hasbeen used to obtain a local 1D model of theattenuation in Vrancea and surrounding areas (Ivan
20 EO 22 E
O24 E
O26 E
O28 E
O30 E
O
44 NO
46 NO
48 NO
Ukraine
Bulgaria
Hun
gary
Republic of
Moldavia
Bla
ck S
ea
100 km
Figure 1 Location of Vran-cea area in Romania.
74 J Seismol (2007) 11:73–85
2003b). Attenuation corrections have been alsoobtained at most K2 stations (Bonjer et al. 2000),including MLR and VRI, which are located in theproximity of the epicentral area (see station codes inFigure 1), but the values are small (less than 0.005 s inabsolute value). However, attenuation corrections atCVO and OZU, located NW of Vrancea, could not beobtained because of the limited number of reliablerecordings at these stations and possible focusingeffects.
On a global scale, the Q factor derived by varioustechniques for the upper mantle is assumed to be inthe range of 120–150 or an average QP of about 200,for the upper 400 km (Sailor and Dziewonski 1978;Anderson and Hart 1978, Bhattacharyya et al. 1996).A similar average of 150 has been derived for theupper mantle beneath the ocean basins by Sipkin andJordan (1980). The values reported by Flanagan and
Wiens (1998) for the Lau back arc basin are in therange of 102–121 for the depth interval of 0–200 km,being associated with the active oceanic subduction inthe Tonga–Fiji area. The highest attenuation (very lowQ values, about 90) has been found within the upper100 km beneath the active portions of the Lau basin byRoth et al. (1999), in the frequency band of 0.1–3.5 Hz.It correlates well with zones of low P-wave velocity. Alocal attenuation tomography indicated both horizontaland vertical variations of Q, the slab being a volumewith higher velocity/low attenuation material (QP>900). Such direct correlation between Q and velocityis supported by other studies conducted in subductionareas (Eberhart-Phillips et al. 2005).
In this paper, the seismic attenuation in the Vranceaarea is quantitatively analysed from global recordingsduring four major (5.9≤Mw≤7.5), intermediate-depth(85≤h≤137 km) events, having near-vertical fault
+
MLR
CVO +
OZU
+
+ VRI
+
+
+
GRE
10 km
+
+
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BUZAU
/ /FOCSANI
TG.SECUIESC
BRASOV
ONESTI
CAPIDAVA-OVIDIU
FAULTTROTUS FAULT
PEC
EN
EA
GA
-CA
ME
NA
FAU
LT1
4
2
3
5
6
7
8
9
a
EAST-EUROPEAN PLATFORM
3
MOESIANPLATFORM
45o
46o
26o
27o
Figure 2 Tectonic settingsin Vrancea area after Hauseret al. (2001). 1: Neogene toQuaternary volcanic rocks;2: Tertiary sedimentaryrocks of the TranssylvanianBasin; 3: Carpathian fore-land and intramontaneousbasins; 4: Neogene piggy–back basins; 5: Subcarpa-thian nappe; 6: marginalfolds nappe; 7: Tarcaunappe; 8: Inner Moldavidenappes; 9: Outer Dacidenappes and Cretaceousposttectonic cover; a: Medi-an Dacide nappes and Cre-taceous posttectonic cover.Local seismological stationsare indicated by triangles.
J Seismol (2007) 11:73–85 75
planes with directions about 230°. As a result,extremely clear recordings of the depth phase pP areobserved especially at a very large number of stationsin North America, with azimuth angles about 320°.Consequently, highly accurate estimations of the QP
factor can be obtained by using the spectral ratiomethod (e.g., Bock and Clements, 1982), especially inthe northwest part of the epicentral area, where thepresence in the upper mantle of a low attenuationvolume is a question of debate in relation to variousgeodynamic models trying to explain Vrancea seis-micity. A very high attenuation area (QP values downto 33) is observed at stations with an azimuth in therange of 306–347°, while for recordings at differentazimuth angles, QP values routinely exceed 200. Verylikely, the low attenuation values are related to anupwelling mantle material located immediately be-neath the crust, but limited in depth to at least100 km; this places limitations on some very recentgeodynamic models which assume that a largevolume of high-temperature, low-attenuation upwell-ing mantle material is present in the northwest ofVrancea (e.g., Chalot-Prat and Gîrbacea 2000).
Methodology
The quality factor QP has been evaluated by using thespectral ratio method, closely following the method-ology described by Roth et al. (1999). It is assumedthat the amplitude spectrum of a certain phasedecreases with the frequency f, in accordance with
A fð Þ ¼ A0 exp �πt*f� �
; ð1Þ
where A0 is a constant depending on the radiationpattern and the attenuation parameter is
t* ¼Z
ray path
dt=Q: ð2Þ
Consider the pair of P and pP phases generated bythe same event and recorded by the same instrument.For deep or intermediate-depth earthquakes, theupgoing longitudinal ray traveling from the hypocen-ter to the free surface, reflecting and traveling on tothe recording station as compressional wave as well,
is denoted as pP phase (Figure 3). At teleseismicdistances, it has a quite similar path to a P wave,except for the area above the hypocenter. Fromequation (1), the logarithm of the amplitude spectraof pP and P phases depends linearly on frequency as
ln App fð Þ.Ap fð Þ ¼ C � π t*
pP� t*p
� �f ;
hð3Þ
where C is a certain constant. In real life, equation (3)is used in a certain frequency range, routinely close to0.2–2 Hz, several basic assumptions being consideredvalid. Thus, the recording sensor is supposed to havea non-zero output in the analysed frequency band. Itis also assumed that for teleseismic distances, there isno frequency dependence of the source radiationpattern, which is a function of the take-off angleonly. The attenuation is related only to the slope ofequation (3), being not depending on the A0 valuescorresponding to the P and pP waves. Finally, thecontamination of the P wave spectrum by noise and ofthe pP spectrum by both noise and P reverberationsare supposed to be negligible both. If so, equation (3)can be used irrespective of the instrument type;however, broadband sensors are more sensitive to
100 km
PpP
pP
NP2
NP1
Surface bounce point
Figure 3 Ray paths of P and pP waves from the 27 October2004 event toward a station at approximately 83° distance.Radiation pattern according to Harvard CMT focal mechanismis also plotted. The cross indicates the surface bounce point ofpP wave. NP1 and Np2 are the corresponding nodal planes.
76 J Seismol (2007) 11:73–85
the micro-seismic noise, when compared to shortperiod seismometers. From equation (3), the attenua-tion factor can be derived as
QP ¼ π tpP � tp� ��
S; ð4Þ
where S is the absolute value of the slope of theregression line fitted to (3). The arrival times of the Pand pP phases are marked as tp and tpP, respectively,being read by hand on seismograms.
Consider a pair represented by an earthquake and arecording station. Three time windows are used toevaluate the corresponding amplitude spectra of P, pPand the reference noise. The start of the P and pPwindows is represented by the corresponding wavearrival. For noise, the end of the time windowanticipates the P wave onset. The windows’ length
depends on the instrumental sampling rate and on theevent magnitude, but is the same for P, pP and thenoise. Routinely, such a window shows a good focusof the wave energy, the vertical and radial traces(where available) being quasi-parallel to each otherfor the entire length. An example is presented inFigure 4, illustrating a window of 6.4 s (128 pointsfor instruments sampling at 20 Hz, or 256 points forsampling at 40 Hz). For instruments at 60 Hzsampling, windows are 8.5333 s in length (i.e., 512points) (Figure 5). Such a time length also avoids thecontamination of the pP wave by sP phase. However,the sP phase could not be clearly identified on almostany of the recordings performed at Canadian and USstations (Figure 6), where the pP phase shows aremarkable high amplitude, routinely exceeding the Pwave amplitude because of the particular radiation
YKW1:BHE
29922.0
-43531.0
67861.0
-114960.0
YKW1:BHN
285735.0
-161030.0
YKW1:BHZ
0.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4
(3)
(2)
(1)
(4)
(5)
(6)
(7)
4.0 5.0 6.03.02.00.0 1.0
TIME [min]
YKW1:BHZ
41791.0
-27296.0
18659
-10024
R
285735.0
-161030.0
1.2156e+05
-72873
YKW1:BHZ
R
P pP
TIME [sec]
Figure 4 Recording of 30 May 1990 event at Yellow Knifearray site YKW1 (STS-1 instrument). Teleseismic P and pParrivals are indicated on the vertical trace 3. Traces 4 and 5 are6.4-s length time windows immediately following the P arrival,
on both vertical (trace 4) and radial (trace 5) channels. Traces 6and 7 are similar windows immediately following the pP onset.Note the quasi-parallelism of the traces 4 and 5, respectively, oftrace 6 and 7.
J Seismol (2007) 11:73–85 77
pattern of the Vrancea earthquakes toward NorthAmerica. That aspect is also minimising the contam-ination of the pP wave by P reverberations. Spectralcomputations for the P, pP and noise windows havebeen performed by using the FFT subroutine de-scribed by Stearn (1975). For almost all the processedwindows, a clear change in the spectral slope about2 Hz is observed (Figure 7), but a sharp limit isdifficult to be assumed. Consequently, the slope of theregression line in equations (3) and (4) has beenevaluated for several frequency bands, such as 0.117–1.992, 0.234–2.109 Hz and so on, where, routinely,the amplitude spectra of both P and pP are clearly
above the noise (Figure 7). Band-passing with a zero-phase Butterworth filter in the range 0.15–2.5 Hz and/or tapering by a 10% cosine taper (Roth et al., 1999)show no significant modification of the Q estimations.For example, the PD02 station (distance Δ=83.4°;azimuth Az=329°) belongs to the Pinedale Array (seeTable 1) and the derived values in relation to the 27October 2004 event are δt*=0.5864 s (i.e., QP=45) inthe frequency range of 0.156–1.875 Hz and δt*=0.7669 s (QP=34) in the range of 0.312–2.031 Hz.These values are just slightly modified for the rest ofthe Pinedale Array stations, irrespective of therecording instrument (short-period or broad band). If
0.0 10.0 20.0 30.0 40.0
(1)
TIME [ sec ]
(2)
(3)
(4)
SNB
:SH
Z
529.0
-610.0
349.0
-343.0
529.0
-610.0
8.0
-1.0
PpP
SNB
:SH
ZSN
B:S
HZ
SNB
:SH
Z
TIME [ sec ]0.0 1.0 2.0 3.0 4.0 5.0 6.0 7.0 8.0
Figure 5 Recording of 30 May 1990 event at Saturna Island(SNB) station (Δ=82°, Az=340°). Trace 1 is the originalseismogram, including noise before P and pP arrivals. Traces 2corresponds to the first 8.53 s of the P phase, while trace 3corresponds to the pP phase. Trace 4 is the noise preceding P
phase arrival. Note the depletion in higher frequencies of pPphase (trace 3) with respect to P phase (trace 2), assumed to bean effect of the attenuation above hypocenter. Recordinginstrument used is Geotech S-13 seismometer.
78 J Seismol (2007) 11:73–85
the above presented band pass filter is applied, the Qvalues in the two frequency bands are 44 and 30,respectively, or 47 and 36 if a 10% cosine taper isused. They are 48 and 39 if both filtering and taperingthe input data. Therefore, no filtering or tapering isused in this study. Finally, the individual 1/Q valuesfor all the stations of the same network in relation to acertain earthquake are averaged. A standard statistical
analysis is subsequently performed, providing 95%confidence error estimates.
Data collection
All digital waveforms available at the IRIS Data Baseand the Geological Survey of Canada networks relatedto the major Romanian earthquakes since the onset ofdigital instrumentation have been investigated for clearP and pP phases and a good signal-to-noise ratio,routinely exceeding 10. Three of the surveyed events(4 March 1997, 30 May 1990, and 27 October 2004)are located at about 90 km in depth, while the 30August 1986 event has its hypocenter at 137 km.Waveforms from the 31 May 1990 event (Mw=6.3,depth=87.3 km) have also been examined but noclear depth phases could be observed because of theevent’s focal mechanism (orthogonal, unlike those ofthe events used in this study). Most of the recordinginstruments have vertical short period sensors, with avariety of constructive types, many of them beingGeotech S-13 and Mark L-4 (c). Instrument responsesare quite similar for almost all SP sensors in the rangeof 0.05–2.55 Hz (Figure 8). Broadband instrumentsare also available (Figure 9). All instruments show a
90o
60o
30o
-60o
-90o
-120o
-150o
Figure 6 North America stations (triangles) and arrays(squares) used in this study.
Q=35
Q=33
6.4923
-4.56946.4923
-4.56946.4923
-4.5694
-3.2512
2.5467
Log
Am
plitu
de
P
pP
noise
pP-P
Frequency [Hz]
0.0 0.5 1.0 1.5 2.5 3.52.0 3.0 4.0
Figure 7 Logarithm of theamplitude spectra for thetime windows from Figure 5(below 4 Hz). DC values arearbitrary. Traces (1)–(3) arethe waves and noise spectra.Trace (4) is the differencepP–P. Note the approxi-mately linear decay below2 Hz, followed by an almostconstant spectrum. Theregression lines for thefrequency window ranges0.117–1.992 and 0.234–2.109 Hz are presented to-gether with the corres-ponding QP values.
J Seismol (2007) 11:73–85 79
Table 1 ISC location of the processed events (with Harvard CMT) and the main attenuation results derived at the various networks/arrays
Event parameters Network Number ofstations
Sensors Δ (°) Az (°) Frequencyband (Hz)
Q
Min Max Min Max
1977-03-04 (Mw=7.5)19:21:54.09 45.8273 N;26.7179 E; 85.8 kmNP1: 50°/28°/86°NP2: 235°/62°/92°
ECTN 3 Geotech S-13 65 66.2 310 312 0.234–1.875 51−11+18
0.469–2.1090.117–1.9920.234–2.109
1986-08-30 (Mw=7.2)21:28:35.97 45.5373 N;26.3138 E; 137.0 kmNP1: 39°/19°/70°NP2: 240°/72°/97°
ECTN 19 S-13 59.3 68.5 306 318 0.234–1.875 117−11+15
0.469–2.1090.117–1.9920.234–2.109
WCTN 17 S-13 80.8 82.6 340 343 0.234–1.875 93−8+8
0.117–1.7580.352–1.992
SEA 71 Mark L-4(c) S-13 81.7 85.4 336 340 0.098–1.759 89−4+4
0.195–1.8571990-05-30 (Mw=6.9)10:40:06.35 45.8474 N;26.6625 E; 89.0 kmNP1: 33°/29°/70°NP2: 236°/63°/101°
ECTN OTT 16 S-13 59.9 68.5 306 315 0.234–1.875 34−2+3
0.469–2.1090.117–1.9920.234–2.109
CLTN 6 S-13 61.5 62 310 311 0.156–1.7190.312–1.875
WCTN 20 S-13 80.6 82.4 340 344 0.234–1.875 43−2+2
0.117–1.9920.469–2.1090.234–2.109
SLC 40 Mark L-4 84.2 89.1 328 331 0.098–1.761 50−3+2
S-13Geotech 18300Episensor 0.196–1.858Benioff 4681aRanger SS-1
SEA Mark L-4(c) 81.5 85.4 331 341 0.099–1.776 46−1+1
S-13 0.197–1.875SS-1
SLCY 9 Mark ProductsL-4
81.8 82.6 331 331 0.098–1.758 59−3+4
0.195–1.8550.195–1.9530.293–2.051
YellowknifeArray (YK)
18 STS-1 67.3 67.5 341 341 0.156–1.875 33−2+2
S-13 0.312–2.0312004-10-27 (Mw=5.9)20:34:36.81 45.79 N;26.62 E; 95.0 kmNP1: 335°/19°/27°NP2: 219°/81°/107°
OTTR 5 S-13 76.9 81.8 341 347 0.195–1.758 48−9+14
WCTN 0.391–1.953POLR_NA 22 Guralp
CMG-3R61.1 66.7 333 344 0.156–1.719 62−6
+7
0.312–1.875YellowknifeArray (YK)
19 Streckeisen 67.4 67.6 342 342 0.156–1.875 57−4+4
STS-1 0.312–2.031S-13
CLTN 11 GuralpCMG-3R
59.7 67.0 310 314 0.195–1.758 100−19+30
ETCN S-13 0.391–1.953
80 J Seismol (2007) 11:73–85
satisfactory recording of the waves’ spectral contentin the range of 0.1–2 Hz (Figure 7, traces 1–3), nocorrelation being observed between the derivedattenuation values and the sensor type. However,broadband instruments are more sensitive to micro-seismic noise when compared to short period sensors;therefore, the signal-to-noise ratio appears routinelyimproved on SP recordings in comparison to thecorresponding broadband ones.
Results
All the stations in Asia or Africa display a near-flatdifferential spectral ratio, with δt* presenting lowvalues, such as 0.10 s, corresponding to a QP of about200. A similar situation is observed at all the stationspertaining to the Eielson Array (see Table 1). For
Table 1 (continued)
Event parameters Network Number ofstations
Sensors Δ (°) Az (°) Frequencyband (Hz)
Q
Min Max Min Max
GAC STS-1POLR_OA 20 Guralp
CMG-3R64.3 72.3 311 323 0.195–1.953 92−11
+14
0.391–2.148SLC 30 Mark L-4 (c) 84.7 88.7 328 331 0.195–1.758 51−4
+4
EpisensorS-13 0.391–1.953Geotech 18300GuralpCMG-3ESP
SEA 45 Mark L-4 (c) 81.5 87.3 337 341 0.195–1.758 52−4+5
S-13 0.391–1.953GRNSNSS-1
SLCY 11 Mark L-4(c) 81.9 82.7 330 331 0.195–1.758 56−7+9
S-13 0.391–1.953GuralpCMG-40T
PinedaleArray (PD)
14 KS-5400 83.4 83.4 329 329 0.156–1.875 38−1+2
Geotech 23900 0.312–2.031Eielson Array(IM)
18 KS-5400 69.4 69.5 357 357 0.156–1.875 >200Geotech 23900 0.312–2.031
Lajitas Array(TX)
5 S-13 92.5 92.5 318 318 0.156–1.875 62−11+17
0.312–2.031
Instrument type and minimum/maximum values for distance and azimuth are also indicated.
OTT: Geological Survey of Canada (GSC) Regional Stations, Ottawa; WCTN: GSC Western Canada Telemetered Network (TN),Sidney, BC; ECTN: GSC Eastern Canada TN, Ottawa; POLR_N(O)A: POLARIS Northern (Ontario) Array; CLTN: GSC CharlevoixLocal TN, Ottawa; SLC: University of Utah, Salt Lake City, UT, USA; SEA: Geophysics Program, University of Washington, Seattle,USA; SLCY: University of Utah Yellowstone, Salt Lake City, UT, USA.
10
9
8
7
6
5
40
Log
Am
plitu
de [
coun
ts]
0.5 1.5 2.51.0 2.0Frequency [Hz]
Geotech 23900
Mark L-4c
Geotech 18300
S-13@1 HzGeotech Benioff 4681a
Ranger SS-1
Figure 8 Instrument response in the range 0.05–2.55 Hz for avariety of short-period sensors used in this study.
J Seismol (2007) 11:73–85 81
example, the IL18 site (Δ=69.4°; Az=357°), belongsto the Eielson Array and the derived values in relationto 27 October 2004 event are δt*=0.1184 s (i.e., QP=202) in the frequency band of 0.156–1.875 Hz, and aδt*=0.1138 s (QP=210) in the range 0.312–2.031 Hz.
In contrast, all the stations with azimuth values inthe range of 306–347° give very low values of QP,irrespective of the event/station pair (Figures 10 and11). The results are quite similar for a particular eventrecorded by stations of the same network/array, thedispersion of the QP values being very low in thiscase (Table 1).
Assuming a certain model with radial symmetryand linear variation of both Q and velocity inside eachlayer, the integral (2) can be evaluated analytically.For the IASPEI91 model with QP values of about1,460 above 210 km depth and a Moho discontinuityat 40 km, the crust contribution to tpP* is of the orderof 0.01 s for a station at 83.4° distance and an event at95 km in depth. That value is just slightly increased if
0 0.5 1.5 2.51.0 2.0Frequency [Hz]
Log
Am
plitu
de [
coun
ts]
12
10
8
6
4
STS-1
GRNSNCMG-40T
CMG-3R
KS-5400
Figure 9 Instrument response in the range 0.05–2.55 Hz for avariety of broadband sensors used in this study.
45o
46o
26o
27o
10 km
117 +15
-11
+
MLR
+
VRI
OZU
+ CVO
+
+
+
PET
FOC
+
SEC
+
GRE
1986-08-30
89+4
-4
93+8
-8
Figure 10 Vrancea epicen-tral area (IRIS catalogue).Squares indicate earth-quakes having Mw≥5.0, inthe depth range 85–161 km.Crosses indicates surfacebounce points of pP wavecorresponding to stationswhere low QP values (QP≤150) have been obtained inrelation to the 30 August1986 event. Filled circlesare surface bouncing pointfor stations with high QP
(QP≥200). Focal mecha-nism is Harvard CMT.Numbers indicate QP valuesat various networks with a95% confidence level.
82 J Seismol (2007) 11:73–85
the local 1D attenuation model (Ivan 2003b) with QP
values in the range of 301–536 is considered bymodifying the IASPEI91 model to above 170 km indepth. Given that no particular features can beobserved on surface tectonic settings northwest of theepicentral area, the crust contribution to the observedabnormal attenuation is negligible, if any. The current-ly available information on Vrancea suggests thepropagation time of teleseismic P wave inside theseismogenic volume to be less than 20 s and an average(reference) value of QP above 200. It follows fromequation (2) that the attenuation inhomogeneity fromthe NW Vrancea volume is most likely located in theupper mantle. Given that the only difference in pathbetween the pP and P phases is in the source area, theobserved lateral differences in Q are most likelyassociated to the pP path above the hypocenter. HighQ values observed at Eielson Array in relation to 27October 2004 event suggest the low-attenuationvolume is limited to the East by the 26°30′ meridian.
Low attenuation (Q>150) observed to some stations(e.g., A16, A21, CBRQ, DELO, EYMN, GAC,KILO) having the lowest azimuth values in relationto the same event could indicate the southeasternboundary of the high attenuation body. However, thesmall number of events available for processing andthe vertical position of their foci does not allow theexact delineating of the low QP volume’s geometry.
Conclusions
A spectral ratio estimate of attenuation has beenperformed in the frequency band of 0.2–2 Hz, inrelation to four major Vrancea events recordedglobally. The physical or structural meaning of thethreshold around 2 Hz is difficult to estimate.However, various attenuation studies performed onPKP waves at distances of about 150° indicated avery similar frequency range of linearity for the
45o
46o
26o
27o
10 km
2004-10-27
1977-03-04
1990-05-30
+
MLR
+
VRI+VRI
+
OZU
CVO
FOC
+
+
GRE
++ PET
+SEC
+
51 +18-11
33 +2-2
34 +3-2
57 +4-4
51 +4-4 / 38 +2
-162 +17
-11
50 +2-3 / 52 +5
-4
/ 56+9-7
100 +30 -19
46+1-1 / 48
+14-9
62 +7-6
/ 43+2-2/ 48
+14
59 +4-3/
92 +14-11
Figure 11 Same caption asFigure 4 in relation to 4March 1977, 30 May 1990,and 27 October 2004events.
J Seismol (2007) 11:73–85 83
spectral ratio, i.e., 0.2–2 Hz. This suggests that theearth’s mantle acts as a band-pass filter in the range of0.2–2 Hz for P waves recorded at teleseismic dis-tances, or that the energy of the P waves is mainlyfocused in the period range between 0.5 and 5 s.
The derived average QP values are about 200, inagreement with various previous results (Sailor andDziewonski 1978; Anderson and Hart 1978; Sipkinand Jordan 1980; Bhattacharyya et al. 1996). Noevidence regarding a possible frequency dependenceof QP in the frequency band 0.2–2 Hz has beenobserved.
A very high attenuation volume (QP values down to33) has been observed for stations located in thenorthwestern part of the Vrancea seismogenic volume,in agreement with the low Q vales obtained byFlanagan and Wiens (1998) and Roth et al. (1999) inthe Lau basin area. No clear correlation of attenuationvalues is observed in relation to shallow geologicalsettings. Most likely, the abnormal Q values are relatedto an upwelling mantle material located immediatelybeneath the crust, but limited in depth to at least100 km, possibly in direct correlation with the smalllow velocity suggested by tomographic studies (Martinet al. 2005). This result places some limitations on thegeodynamic model of Chalot-Prat and Gîrbacea(2000), which assumes a large volume of high-temperature, high-attenuation upwelling asthenospherematerial to be present northwest of Vrancea.
Acknowledgement The author is thankful to Rick Benson,Anh Ngo and to the IRIS DMC team for support in acquiringthe digital waveforms. Also to the people maintaining the EventWaveform Archive of the Geological Survey of Canada (http://seismo.nrcan.gc.ca/nwfa/events/). Tongji University, School ofOcean and Earth Sciences is acknowledged for the excellentresearch conditions provided during a Visiting Professor stage.Comments and various suggestions of two anonymous refereesdefinitely improved the paper. GMT files (Wessel and Smith1996) were used to prepare some of the diagrams.
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