Geophys. J. R. astr. SOC. (1974) 36,227-233

Research Note

Earthquakes that look like Explosions A. Douglas, J. A. Hudson, P. D. Marshall and J. B. Young

(Received 1973 July 11)

The mb : h4, criterion is now generally accepted as the best method of distinguishing on seismic evidence between earthquakes and underground explosions (see for example SIPRI, 1968; Marshall & Basham 1972): for explosions the body wavemagnitude (&) is usually 1-2 magnitude units greater than the surface wave magnitude (Ad,); for earthquakes mb and M , are more nearly equal. Some earthquakes (known to be earthquakes because of their focal depth: see below) however fail on the mb: M , criterion (Landers 1972; CCD 1972) and in this paper we suggest reasons why these earthquakes fail and also why paradoxically these earthquakes may turn out to be the seismic events that are easiest to identify.

Fig. 1 (a) shows the P signals from such an earthquake with epicentre near Alma Ata, Kazakhstan, as recorded at Yellowknife, Canada (YKA). The United States Coast and Geodetic Survey (USCGS) give an average mb for this earthquake of 4.9 (mb measured at YKA is 5*0), M, measured, using the methods of Marshall & Basham (1972), from the World Wide Standard Station Network (WWSSN) is 3.7. On the nib : M, criterion using the results of Marshall & Basham (1972) this earthquake lies close to the explosion population. The signal consists of two clear arrivals (which is typical of this type of anomalous earthquake) and at first sight might appear to in- dicate two explosions fired six seconds apart. Closer examination of the signal however shows the second arrival is of opposite polarity to the first and is thus probably the surface reflection p P . Confirmation that the second pulse is of reversed polarity is given in Fig. 1 (b) where the effects of anelastic attenuation and recording instrument have been removed from the recorded signal using the spike filtering technique described by Douglas et al. (1972b). The record now becomes simply a positive impulse (P) followed by a negative impulse (pP) . The P-pP time appears to be about 6 s indicating a depth of focus of about 20km. The event is thus con- clusively identified as an earthquake.

To construct a spike filter some value has to be assumed for t* where f * = T J Q A v ;

T is the total travel time and QAV is the average quality factor Q, along the path Alma Ata to YKA. T is well known from travel-time studies so that in assuming a vdue oft* we have effectively to assume a value for Q A p If the correct value of QAV

is used the spiked seisniogram will show the shape of the P and p P pulses (and any other prominent pulse) at source. If now we assume these pulses at source are either wholly positive or wholly negative (i.e. they do not oscillate about the zero baseline) we have a way of estimating QAv: we simply spike the record with a range of values of t* and select the value that gives a spiked seismogram with pulses that do not oscillate about the baseline. (To obtain pulses that oscillate requires that the earth- quake source oscillates and the theoretical studies of Burridge (1968) using an earth- quake model of slip on a plane shows that any oscillations are likely to be negligible where friction resists slipping.) From Fig. 1 for example we see that using a value of t* = 0.4 gives P and p P pulses that swing strongly positive and negative whereas for I* = 0.2, P and p P are almost simple impulses. Even with f * = 0.2 however the P pulse seems to have a small overshoot. Using smaller values of f * reduces the

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228 A. Douglas et ai.

FIG. 1. P signal from the Alma Ata earthquake of 1968 July 1 as recorded at Yellowknife, Canada together with spiked seismograms. (a) recorded signal,

(b) spiked seismogram using t* = 0 . 2 , (c) spiked seismogram using t* = 0.4.

overshoot slightly but as we will show later there is another explanation of this overshoot.

Thus it would appear that t* = 0.2 is about the value for the Alma Ata-YKA path and thus that QAV is about 3500. Now this value of QAV is very high: Teng (1968) gives values of QAV of around 200 and the largest value of Q measured so far appears to be 1700, (Frasier & Filson 1972). The very high value of Qnv obtained here suggests an explanation of why some earthquakes fail on the mb: M , criterion. Ward & Toksoz (1971) have pointed out that the differences between mb and M , tend to be larger for events in Central Asia than for North America and suggest that at least part of this regional dependence is because body waves are attenuated less (Q is higher) beneath Central Asia than beneath North America; for a given M,, mb is larger for events in Central Asia than in North America. However, when comparing earthquakes from Central Asia with explosions from North America, Marshall & Basham (1972) found no overlap of earthquakes and explosions on the m,, : M , criterion for the group of events they studied. Anomalous earthquakes of the type discussed here may be those that are observed on paths of almost no attenuation and so they do overlap with the explosion population. Note also that no sP phase is obvious in Fig. l(a). A signal of this type consisting only of P and p P will be radiated to long range by a 45" dip slip fault. Douglas, Hudson & Kembhavi (1971b) and Douglas, Hudson & Blamey (1972a) have shown that it is just this orientation of fault plane that produces large mb relative to M,.

To support these explanations of why this Alma Ata earthquake looks like an explosion on the mb : M , criterion we show in Fig. 2 (a) the P signal computed using

Research Note 229

V V 0 L - - - L - l s

FIG. 2. Computed P signals obtained from proposed model of Alma Ata earthquake (a) computed seismogram including effects of attenuation and recording instrument, (b) computed seismogram spiked using f * = 0.2, (c) computed seismogram spiked with t* = 0.4, (d) computed seismogram excluding the effects of attenuation

and recording instrument.

the methods of Douglas et al. (1972a) for a 50" dip slip fault with the observer on an azimuth 90" to the strike of the fault plane and at a distance of 80" (YKA is 73.3" from Alma Ata) and with t* = 0.2. The agreement between the computed signal and observed is clearly very good. The details of the source model and crustal structures used are given in Tables 1 and 2, respectively. We arrived at this model by a process of trial and error starting from a 45" dip slip fault and adjusting various parameters of source structure and fault dimensions and orientation until an acceptable fit was obtained with observation. We also obtained a clue as to what source models to try from the form of the P and p P pulses in Fig. l(b). These pulses are almost impulses but not quite: p P in particular shows a trailing edge that is sharper than the leading edge; for P the reverse appears to be true but this is less easy to see and there is the added complication of an apparent overshoot. If we assume that the mechanism of

230 A. Douglas et al.

Table 1 Parameters of theoretical earthquake model

Depth of focus = 19.4km Dip of fault plane = 50"

= 90" to strike = 1.25 km

Direction of slip on fault plane Radius of fault plane Fracture velocity (0.6 /3 velocity where = 2.1 km s-l

/3 is S velocity on source layer)

Table 2 P-wave vefocity S-wave velocity Density Thickness

(km s-I) (kms-I) ( g ~ m - ~ ) (km)

Layer 1 4.80 - 2.7 1.0

Layer 3 8.2 - 3.3 co

Layer 1 6.14 - 2.80 5.3

Layer 3 8.09 - 3.28 co

(a) Source structure

Layer 2 6.15 3.5 2.8 22.4

(b) Station structure

Layer 2 1.28 - 3.20 19.7

Where S-wave velocity (j3) is not listed it is assumed that j3 = s/\/3 where a is the P-wave velocity.

the Alma Ata earthquake is slip on a plane dipping at about 45", most of these differences in pulse shape between P and p P can be understood as the effects of radiation from a moving source of the type propos2.d by Savage (1 966). Fig. 3 shows a vertical section AOB through such a plane at right angles to the strike of the plane. The fault model of Savage (1966) for a disc-shaped fault plane assumes that faulting is initiated at 0 and spreads with uniform velocity in all directions in the plane. After faulting a disc-shaped fault plane has formed; A and B are the edges of the disc. Both AOB and COD (thc auxiliary plane) are nodes of P radiation. Imagine now an observer looking at the source along P,O then as the fracture sprezds from 0 to A there is a component velocity towards the observer which has the effect of steepening the leading edge of the P pulse. When fracturing ceases at A, energy apparently continues to be radiated froin the fracture spreading from 0 to B away from the observer, because of the greater distance this energy has to travel to the observer. The tail of the pulse is thus spread out. There is thus a Doppler shift in the frequency of the radiated signal depending on whether the fracture is approaching or receding from the observer.

For an observer looking at the source along P,O any Doppler effects are small because the plane of fracture is almost at right anglcs to P,O. Here the trailing edge is sharp because the observer at P , sees fracture cease at virtually the same time all round the edge of the fault plane; radiation then ceases abruptly. The leading edge of this pulse is less steep than the trailing edge because the length of the fracture contour is zero at initiation and the amplitude of the radiated pulse only grows as fast as the contour of fracture grows. (In the above discussion we have assumed constant slip over the whole fault plane. The effects are less marked if the more realistic assumption is made as we habe done in computing the seismogram Fig. 2 (a) that the amount of slip falls off towards the edge of the fault.)

The observed differences in pulse shape of P and p P might then be explained by assuming that P leaves the source roughly along O P , and p P roughly along OP, ( p P is reflected at the free surface and follows the P path to the observer) and we used this as a further constraint on the choice of model.

Research Note 23 1

I 'n \ Y

, 4

FIG. 3. Radiation pattern for a 45" dip slip fault model.

Fig. 2 shows in addition to the computed seismogram the spiked signals for t* = 0.2 and t* = 0.4 for comparison with the spiked signals shown in Fig. 1. Note that when the incorrect value oft* is used (i.e. t* = 0-4) for the spiking filter the same kind of overshoot is shown as for the observed signal. Fig. 2(d) shows that the com- puted signal without any effects of attenuation or recording instrument. Comparing Fig. 2(b) and 2(d) shows that using the correct spiking filter does remove most of the effects of attenuation and recording instrument.

Comparing Figs 2(b) and 1 (b) shows that the observed differences in pulse shape between P and p P and the apparent overshoot of the P pulse are reproduced on the computed signal. In the computed signal this apparent overshoot is due to the arrival of a negative pulse generated by the conversion of part of the S energy radiated downwards at the source, to P at the boundary between layer 2 and layer 3 of the source model. This may also be the explanation of the apparent overshoot of the P pulse on the observed signal. Note also that on the computed signal a small reflected pulse from the layer 2 and layer 1 boundary of the source structure arrives just ahead of and overlaps with p P making the p P pulse appear wider than it is and there is some evidence in Fig. I@) that the leading edge of the p P pulse on the observed signal is similarly influenced by an earlier arrival.

(The next step in the analysis of this earthquake would be to try and model the signals from this earthquake recorded at other arrays. Unfortunately the record from Warramunga, Australia is obscured by the signal from another earthquake and the Gauribidanur (GBA), India and Eskdalemuir, Scotland records are almost

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obscured by noise. However the GBA signal does show, as one would expect, two arrivals with about the correct separation for P and pP but the signal/noise ratio is not large enough to allow any conclusions to be drawn about pulse shapes.)

Turning now to the mb : M , differences: the observed difference, mb-M, is 1.3 magnitude units, and for the theoretical model m b - M s is 1.2 magnitude units (assuming a stress drop of 10 bars, mb = 4.9 and M , = 3.7). Agreement between theory and observation is thus excellent.

We have been able to analyse in detail the single P signal studied here very success- fully. The principal reason for this is that the signal was observed on a very high Q path so that any attenuation effects are small and these effects are easily removed by spiking. Also Douglas, Marshall & Corbishley (1971a) have suggested that it is just such signals recorded on high Q paths that should be very simple, any scattered signals being of negligible amplitude. Consequently the source functions of the earthquakes can be readily obtained. If all earthquakes that look like explosions turn out to be those that are recorded on high Q paths then their P signals also should be susceptible to the detailed analysis of the kind described here. From the detailed knowledge of the seismic source obtained from such an analysis it should be possible to identify these events unambiguously as earthquakes.

A. Douglas et a 1.

A. Douglas, P. D. Marshall and J. B. Young: J. A. Hudson:

MOD ( P E ) Dept of Applied Mathematics

University of Cambridge, Blacknest, Brimpton, and Theoretical Physics

Reading RG7 4RS, Berkshire Silver Street, Cambridge

References

Burridge, R., 1968. Theoretical seismic sources and propagating brittle cracks, J. Phy. Earth. 16, Special Issue, 83-92.

CCD 1972. A review of currentprogress andproblems in seismic verijication, Conference of the Committee on Disarmament, CCD /388.

Douglas, A., Corbishley, D. J., Blarney, C. & Marshall, P. D., 1972b. Estimating the firing depth of underground explosions, Nature, 237,26-28.

Douglas, A., Hudson, J. A. & Blarney, C., 1972a. A quantitative evaluation of seismic signals at teleseismic distances-111. Computed P and Rayleigh wave seismograms, Geophys. J. R. astr. SOC., 28,385-410.

Douglas, A., Hudson, J. A. & Kembhavi, V. K., 1971b. The relative excitation of seismic surface and body waves by point sources, Geophys. J. R. astr. SOC., 23, 451-460.

Douglas, A,, Marshall, P. D. & Corbishley, D. J., 1971a. Absorption and the com- plexity of P signals, Nature Phys. Sci., 233,50-51.

Frasier, C. W. & Filson, J., 1972. A direct measurement of the Earth’s short period attenuation along a teleseismic ray path, J. geophys. Res., 77,3782-3787.

Landers, T., 1972. Some interesting central Asian events on the M, : mb diagram, in Seismic Discrimination, Semiannual Technical Summary, Lincoln Laboratory, Massachusetts, June 1972.

Marshall, P. D. & Basham, P., 1972. Discrimination between earthquakes and underground explosions employing an improved M, scale, Geophys. J. R . astr. Soc., 28, 431458.

Savage, J. C., 1966. Radiation from a realistic model of faulting. Bull. seism. SOC. Am., 56, 577-592.

Research Note 233

SIPRI, 1968. Seismic methods of monitoring underground explosions, International

Teng, T. L., 1968. Attenuation of body waves and the Q structure of the mantle,

Ward, R. W. & Toksoz, M. N., 1971. Causes of regional variation of magnitudes,

Institute for Peace and Conflict Research, Stockholm.

J. geophys. Res., 73, 2195-2208.

Bull. seism. SOC. Am., 61,649-670.