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2The Nature of Earthquakes

2.1 Introduction

An earthquake is a spasm of ground shaking caused by a sudden release of energyin the earth’s lithosphere (i.e. the crust plus part of the upper mantle). This energyarises mainly from stresses built up during tectonic processes, which consist of inter-action between the crust and the interior of the earth. In some parts of the world,earthquakes are associated with volcanic activity. For example, in Guatemala suchearthquakes occur in swarms, with an average duration of three to four months, thelargest having a magnitude normally under 6.5. These events are of shallow focusand cause considerable damage within a radius of about 30 km from the epicentre.Human activity also sometimes modifies crustal stresses enough to trigger small oreven moderate earthquakes, such as the swarms of minor tremors resulting frommining in the Midlands of England, or the sometimes larger events induced by theimpounding of large amounts of water behind dams, such as the earthquakes associ-ated with the construction of the Koyna dam in central India in 1967 (Chopra andChakrabarti, 1973).

While the design provisions of this book apply to all earthquakes regardless of origin,any discussion of earthquakes themselves is generally confined to events derived fromthe main cause of seismicity, i.e. tectonic activity.

As most earthquakes arise from stress build-up due to deformation of the earth’scrust, understanding of seismicity depends heavily on aspects of geology, which is thescience of the earth’s crust, and also calls upon knowledge of the physics of the earthas a whole, i.e. geophysics. The particular aspect of geology which sheds most light onthe source of earthquakes is tectonics, which concerns the structure and deformationsof the crust and the processes which accompany it; the relevant aspect of tectonics isnow often referred to as seismotectonics.

Geology tells us the overall underlying level of seismic hazard which may differfrom the available evidence of historical seismicity, notably in areas experiencingpresent day quiescent periods.

Earthquake Risk Reduction D.J. Dowrick 2003 John Wiley & Sons, Ltd ISBN: 0-471-49688-X (HB)

16 The nature of earthquakes

2.2 Global Seismotectonics

On a global scale, the present-day seismicity pattern of the world is illustrated ingeneral terms by the seismic events plotted in Figure 2.1. Most of these events can beseen to follow clearly defined belts which form a map of the boundaries of segments ofthe earth’s crust known as tectonic plates. This may be seen by comparing Figure 2.1with Figure 2.2, which is a world map of the main tectonic plates taken from the highlyunderstandable book on the theory of continental drift by Stevens (1980). Accordingto the latter, the earth’s crust is composed of at least 15 virtually undistorted plates oflithosphere. The lithosphere moves differentially on the weaker asthenosphere whichstarts at the Low-Velocity Layer in the Upper Mantle at a depth of about 50 km.Boundaries of plates are of four principal types;

(1) Divergent zones, where new plate material is added from the interior of the earth.(2) Subduction zones, where plates converge and the under-thrusting one is con-

sumed.(3) Collision zones, former subduction zones where continents riding on plates are

colliding.(4) Transform faults, where two plates are simply gliding past one another, with no

addition or destruction of plate material.

Almost all the earthquake, volcanic and mountain-building activity which marks theactive zones of the earth’s crust closely follows the plate boundaries, and is related tomovements between them.

Divergent boundaries are found at the oceanic sea-floor ridges, affecting scatteredislands of volcanic origin, such as Iceland and Tristan da Cunha, which are locatedon these ridges. As these zones involve lower stress levels, they generate somewhatsmaller earthquakes than the other types of plate boundary.

As can be seen in Figure 2.2, subduction zones occur in various highly populatedregions, notably Japan and the western side of Central and South America. Figure 2.3shows the cross-section of the likely structure of the subduction zone formed by thePacific plate thrusting under the Indian-Australian plate beneath the North Island ofNew Zealand. The seismic cross-section corresponding to Figure 2.3 is in Figure 2.4,and gives earthquakes located under the shaded region of the key map (during a periodof time when no events shallower than c. 40 km occurred). The zone of diffuse seismicactivity which exists down to a depth of over 300 km is believed to be related bothto volcanic activity (movement of magma in the crust and upper mantle and relatedexpansion and contraction), and to faulting within the volcanic belt and the ‘NewZealand Shear Belt’. The latter is a continuation of the major Alpine Fault of theSouthern Alps (Figure 2.5). Below 100 km and down to about 250 km, the pattern ofearthquakes tends to lie on a well-defined plane known as a Benioff Zone, dipping50 degrees to the north-west. This is the contact plane between the Indian-Australianplate and the Pacific plate. The isolated group of earthquakes about 600 km deep inFigure 2.4 have been conjectured to be caused by a piece of lithosphere that has becomedetached and has moved deeper into the mantle, as illustrated in Figure 2.3.

The progressive movement of the Pacific plate, subducting under the overlyingplate, caused shear stresses to develop, as illustrated by Walcott’s (1981) geophysical

Global seismotectonics 17

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Figure 2.2 Tectonic plate map of the world, showing names of the seven largest plates andindicating subduction zones and the directions of plate movement (reproduced withpermission from G.R. Stevens, 1980)

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Figure 2.3 The likely structure of the subduction zone beneath the North Island ofNew Zealand inferred from figure 2.4 (reproduced with permission fromG.R. Stevens, 1980)

Global seismotectonics 19

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Figure 2.4 Seismic cross-section through the North Island of New Zealand, showing locationsof earthquake foci (reproduced with permission from G.R. Stevens, 1980)

study (Figure 2.5), which relates the shear strains of the shear belt referred to aboveto large historical earthquakes. It is believed that the fault forming the plate boundaryperiodically locks together, and this leads to an accumulation of shear and compres-sional strain until it is in part relieved by a large thrust type of earthquake. The suddenrelease of strain (when the shear resistance is overcome) signals the recommencementof movement of the subducting plate in a further cycle of aseismic slip, then anotherlocking of the fault leads to the next plate interface earthquake.

As well as the 15 or so main plates shown in Figure 2.2, studies of seismic activityneed to consider the smaller buffer plates or sub-plates which in certain areas tendto ease the relative movements of the world’s giant plates. Buffer plates have beenrecognized in Tibet and China, the western USA, and at the complex junction of theAfrican, Arabian, Iranian, and Eurasian plates, where eight Mediterranean buffer plateshave been identified (Stevens, 1980).

In the foregoing discussion, tectonic plates have been described as rigid, virtuallyundistorted plates and the world’s principal zones of seismicity have been shown tobe associated with the interaction between the plates. However, occasional damagingintra-plate earthquakes also occur, well within the interior of the plates that clearlyare not associated with plate boundary conditions, and so far their origins are ill-understood. The uncertainties associated with intra-plate seismicity are much greaterthan is the case for interplate regions of high seismicity.


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Figure 2.5 Generalized shear strain rates and large (MW ≥ 6.8) shallow (hC < 50 km) NewZealand earthquakes from 1840–2000 (adapted from Walcott, 1981)

2.3 The Strength of Earthquakes—Magnitude and Intensity

During earthquakes the release of crustal stresses is believed generally to involvethe fracturing of the rock along a plane which passes through the point of origin(the hypocentre or focus) of the event (Figure 4.25). Sometimes, especially in largershallower earthquakes, this rupture plane, called a fault, breaks through to the groundsurface, where it is known as a fault trace (Figure 4.41).

The strength of earthquakes—magnitude and intensity 21

The cause and nature of earthquakes is the subject of study of the science of seis-mology, and further background may be obtained from the books by Richter (1958),Bolt (1999) and Lay and Wallace (1995).

Unfortunately, for non-seismologists at least, understanding the general literaturerelated to earthquakes is impeded by the difficulty of finding precise definitions offundamental seismological terms. For assistance in the use of this book, definitions ofsome basic terms are set out below. Further definitions may be found elsewhere in thisbook or in the references given above.

The strength of an earthquake is not an official technical term, but is used in thenormal language sense of ‘How strong was that earthquake?’ Earthquake strengthis defined in two ways: first the strength of shaking at any given place (called theintensity) and second, the total strength (or size) of the event itself (called magnitude,seismic moment, or moment magnitude). These entities are described below.

Intensity is a qualitative or quantitative measure of the severity of seismic groundmotion at a specific site. Over the years, various subjective scales of what is often calledfelt intensity have been devised, notably the European Macroseismic and the Mercalliscale, which are very similar. The most widely used in the English speaking worldis the Modified Mercalli scale (commonly denoted MM), which has twelve gradesdenoted by Roman numerals I–XII. A detailed description of this intensity scale isgiven in Appendix A, taken from Dowrick (1996).

Quantitative instrumental measures of intensity include engineering parameters suchas peak ground acceleration, peak ground velocity, the Housner spectral intensity,and response spectra in general. Because of the high variability of both subjectiveand instrumental scales, the correlation between these two approaches to describingintensity is inherently weak (Figure 4.23).

Magnitude is a quantitative measure of the size of an earthquake, related indirectlyto the energy released, which is independent of the place of observation. It is calculatedfrom amplitude measurements on seismograms, and is on a logarithmic scale expressedin ordinary numbers and decimals. Unfortunately several magnitude scales exist, ofwhich the four most common ones are described here (ML, MS , Mb and MW ).

The most commonly used magnitude scale is that devised by and named afterRichter, and is denoted M or ML. It is defined as

ML = log A − log A0 (2.1)

where A is the maximum recorded trace amplitude for a given earthquake at a givendistance as written by a Wood–Anderson instrument, and A0 is that for a particularearthquake selected as standard.

The Wood–Anderson seismograph ceases to be useful for shocks at distances beyondabout 1000 km, and hence Richter magnitude is now more precisely called local mag-nitude (ML) to distinguish it from magnitude measured in the same way but fromrecordings on long-period instruments, which are suitable for more distant events. Whenthese latter magnitudes are measured from surface wave impulses they are denoted byMS . Gutenburg proposed what he called ‘unified magnitude’, denoted m or mb, whichis dependent on body waves, and is now generally named body wave magnitude (mb).This magnitude scale is particularly appropriate for events with a focal depth greaterthan c. 45 km. All three scales ML, mb and MS suffer from saturation at higher values.


The most reliable and generally preferred magnitude scale is moment magnitude,MW . This is derived from seismic moment, M0, which measures the size of anearthquake directly from the energy released, Wyss and Brune (1968), through theexpression

M0 = µAD (2.2)

where µ is the shear modulus of the medium (and is usually taken as 3 × 1010 Nm), A

is the area of the dislocation or fault surface, and D is the average displacement or slipon that surface. Seismic moment is a modern alternative to magnitude, which avoidsthe shortcomings of the latter but is not so readily determined. Up to 1985, seismicmoment had generally only been used by seismologists.

Moment magnitude is a relatively recent magnitude scale from Kanamori (1977)and Hanks and Kanamori (1979), which overcomes the above-mentioned saturationproblem of other magnitude scales by incorporating seismic moment into its definition,such that moment magnitude

MW = 2

3log M0 − 6.03 (M0 in Nm) (2.3)

Local magnitude ML is inherently a poor magnitude scale, as shown by the plotin Figure 2.6 of against of New Zealand data from Dowrick and Rhoades (1998),who found that the best fit relationship for estimating MW from ML and depth,hc, was

MW = 0.96[±0.49] + 0.84(±0.08]ML − 0.0055[±0.0015](hc − 25) (2.4)

The regression explains only 59% of the variance and has a residual standard error of0.31. The ML scale as estimated in other parts of the world, as well as New Zealand,is similarly unreliable.

The relation between moment magnitude MW , surface-wave magnitude MS andcentroid depth hc, using data restricted to modern MW determinations (i.e. from 1964March 8 onwards), is shown in Figure 2.7. For earthquakes of hc ≤ 30 km, MS and MW

are close to being equal above magnitude 6.5. At lower magnitudes MS is consistentlysmaller than MW , and is as much as a quarter-unit smaller between magnitude 5.0and 5.5. Depth also influences the discrepancy between MS and MW ; for deep NewZealand earthquakes (hc > 50 km) MS is about a half-unit smaller than MW betweenmagnitude 5.0 and 5.5. This results from the tendency for MS to decrease with depthfor earthquakes of a given seismic moment. Karnik (1969) first dealt with this effectby proposing a focal depth correction term for MS in relation to mb for various partsof Europe, while Ambraseys and Free (1997) more recently estimated a focal depthcorrection term for in relation to log M0 for European earthquakes. Considering theNew Zealand data (Figure 2.7) Dowrick and Rhoades (1998) found the best fit forfinding MW in terms of MS and hc was the quadratic expression:

MW = 1.27[±0.16] + 0.80(±0.03]MS + 0.087[±0.031](MS − 6)2 + 0.0031

× [±0.0006](hc − 25) (2.5)

The strength of earthquakes—magnitude and intensity 23

Figure 2.6 Scatter plot of local magnitude ML against moment magnitude MW for NewZealand earthquakes distinguishing events in different classes of centroid depthhC . Also shown are the linear and quadratic regression fits for ML evaluated athC = 25 km and a local regression trend curve of ML on MS for events withhC ≤ 50 km (from Dowrick and Rhoades, 1998)

The above expression explains 93% of the variance. In Eq. (2.5) it can be seen that thequadratic term contributes significantly to the regression because the coefficient of thisterm is more than twice its standard error. It is of interest to note that, although theirexpression is different from ours, Ambraseys and Free obtained a coefficient for theirdepth term of 0.0036, which is very similar to the coefficient of 0.003 in Eq. (2.5).

Also shown in Figure 2.7 is the relation of Ekstrom and Dziewonski (1988), derivedfrom global data, between log MO and MS for events with h < 50 km. In terms of MW ,this relation is

MW =

2.13 + 23MS MS < 5.3

9.40 − √41.09 − 5.07MS 5.3 ≤ MS ≤ 6.8

0.03 + MS MS > 6.8

(2.6)

As seen in Figure 2.7, there is no great difference between this relation and the linearand quadratic fits for shallow New Zealand events over the magnitude range of thedata, but the latter also describe the effect of depth.


Figure 2.7 Scatter plot of moment magnitude MW against surface wave magnitude MS forearthquakes distinguishing events in different classes of centroid depth hC . Alsoshown are the linear and quadratic (Eq. (2.5)) regression fits for MW evaluated athC = 25 km and a local regression trend curve of MW on MS for events with hC ≤50 km, and the relation of Eckstrom and Dziewonski (Eq. (2.4) (from Dowrick andRhoades, 1998)

References

Ambraseys NN and Free MW (1997) Surface-wave magnitude calibration for European regionearthquakes. J Earthq Eng 1(1): 1–22.

Barazangi M and Dorman J (1969) World seismicity map of ESSA coast and geodetic surveyepicentre data for 1961–67. Bull Seism Soc Amer 59: 369–80.

Bolt B (1999) Earthquakes. WH Freeman & Co, New York, 4th Edn.Chopra AK and Chakrabarti P (1973) The Koyna earthquake and damage to Koyna dam. Bull

Seism Soc Amer 63(2): 381–97.Dowrick DJ (1996) The Modified Mercalli earthquake intensity scale—Revisions arising from

recent studies of New Zealand earthquakes. Bull NZ Nat Soc Earthq Eng 29(2): 92–106.Dowrick DJ and Rhoades DA (1998) Magnitudes of New Zealand earthquakes, 1901–1993.

Bull NZ Soc Earthq Eng 31(4): 260–80.Ekstrom G and Dziewonski AM (1988) Evidence of bias in estimations of earthquake size.

Nature 332: 319–323.Hanks TC and Kanamori H (1979) A moment magnitude scale. J Geophys Res B 84: 2348–50.Kanamori H (1977) The energy release in great earthquakes. J Geophys Res 82: 2981–87.

References 25

Karnik V (1969) Seismicity of the European Area, Part 1. Reidel, Dordrecht, Holland.Lay T and Wallace TC (1995) Modern Global Seismology. Academic Press, San Diego.Richter CF (1958) Elementary Seismology. Freeman, San Francisco.Stevens GR (1980) New Zealand Adrift. AH & AW Reed, Wellington.Walcott RI (1981) The gates of stress and strain. In: Large Earthquakes in New Zealand. The

Royal Society of New Zealand, Miscellaneous Series No 5.Wyss M and Brune J (1968) Seismic moment, stress and source dimensions. J Geophys Res 73:

4681–94.

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