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CHAPTER 1

NATURE OF EARTHQUAKES

1.1 Living Earth: continental drift

Figure 1.1 shows that earthquake activity is mainly confined to relatively narrowbands of intense seismicity. Earthquakes occurring along these bands are interplate

earthquakes. The most intensive concentration of earthquake activity is along a belt

around the circumference of the Pacific Ocean along the western coasts of South,

Central and North America, along the Kurile islands, the Aleutian arc, Japan,

Indonesia and the Philippines, Papua New Guinea and New Zealand. Many active

volcanoes also affect this area. Approximately three-quarters of the global seismic

activity occurs in this region. Another region of intense earthquake activity is in the

region of the Azores islands in the Atlantic Ocean and extends across the

Mediterranean Sea, affecting Portugal, Spain and the countries of the Mahgreb, Italy,

Greece, Turkey, the Levant, Iran, Pakistan, northern India, Nepal and China. Seismic

activity in this region is about one fifth of the global activity. Intensive earthquakeactivity can also be observed throughout the oceans of the world, including the middle

of the Atlantic, the south and southeast of the Pacific, and the Indian oceans. These

mid-ocean earthquakes actually can serve for explaining the mechanism of the

earthquake patterns described above.

Figure 1.1. Earthquake activity around the world in the period from 1977 to 1994

(Shearer, 1999).

Figure 1.1 also shows that few earthquakes are far from the major zones of seismicactivity and they are located in Australia, eastern North America, central India and

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northeast Brazil (stable continental regions). These earthquakes are called intraplate

events and their explanation for their occurrence is not as straightforward as that for

the interplate earthquakes. The number of intraplate earthquakes is small but they can

be significant in size. The largest three earthquakes (with known magnitudes)

occurred in the United States of America were intraplate events whose magnitudes are

greater than 8. These earthquakes occurred in the region of New Madrid (Figure 1.2),one of the well-known stable continental regions in the world.

Figure 1.2. Area affected by the New Madrid earthquakes of 1811-12

in the central United States (Algermissen, 1983).

The internal structure of the Earth is one of the key parameters to understand the

major seismic activity around the world. The Earth may be considered to have three

concentric layers (Figure 1.3). The innermost part of the Earth is the core that ismainly composed of iron. The core has two separate parts, the inner and outer cores.

The mantle is between the crust (outermost layer of the earth) and the core. The

abrupt changes in density that are also reflected in the velocity of propagation of

seismic waves (middle and lower panels in Figure 1.3) differentiate the mantle, the

outer core and the inner core. The sudden variation in the density close to the surface

is called as the Moho discontinuity (identified by the Croatian seismologist

Mohorovic) and it is accepted as the boundary between the mantle and the crust. The

crust thickness is 7 km under the oceans and its average thickness is 30 km under the

continents (attains greater values under the mountain ranges). The crust has basaltic

structure under the oceans and it is mainly comprised of basalt and granite under the

continents.

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Figure 1.4. Cross-section of the outermost part of the Earth illustrating the structure of the

lithosphere and asthenosphere and the principle of isostasy

(Press & Siever, 1986).

The interior of the Earth is in constant motion that is driven by heat. The source of

this heat is radioactivity within the core. The temperature gradient across the Earth

sets up a flow of heat towards the surface; in the outer core, which is liquid and the

mechanism of heat transfer is convection. The convection currents within the

astenosphere give rise to movement of the lithosphere (Figure 1.5) that explains the

constant motion of Earth’s interior . The movement within the mantle (i.e., the

astenosphere) creates the movement of the lithosphere; the convection currents simply

carry the lithospheric plates (tectonic plates) rather like a conveyor belt. The

movement of these plates results in two slabs diverging from each other or their

collision; one slab descending beneath the other one.

Figure 1.5. Schematic illustration of heat convection within the Earth’s mantle anda comparison with convection currents in a heated liquid

(Press & Siever, 1986).

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The above physical process is the rational behind the continuous motion of the

continents as well. In fact, 200 million years ago all of the continents had formed a

single landmass, called Pangaea. This continent broke up, initially forming two

continents, Laurasia and Gondwanaland, about 150 million years ago. By 100 million

years ago, Laurasia had split into the continents of North America and Eurasia, and

Gondwanaland had divided into the continents of India, South America, Africa,Antarctica and Australia (Figure 1.6). These continents have continued to move to

bring about their current configuration, including the collision of India with Eurasia

about 50 million years ago.

Figure 1.6. Configuration of the continents about 100 million years before the present

(Tarbuck & Lutgens, 1997).

The key explanations about how the continents had come to move apart was brought

by some pioneering geologists such as Professor Richard Field from the University of

Princeton who saw the importance of studying the geology of the ocean floor. The

mountain chains along the major oceans (oceanic ridges) were found to exist in offsetparallel segments and a very deep gorge ran down the middle of each part of the range

(Figure 1.7). Simultaneously, seismologists, after starting to locate earthquakes by

using the recordings of earthquake waves, found that the location of many

earthquakes coincided with the oceanic ridges, which suggested that these were zones

of movement. In 1960, Harry Hess, a student of Richard Field, put forward the theory

of sea-floor spreading and suggested that material was rising up from within the

mantle of the Earth into the central gorges of the oceanic ridges and forming new

oceanic crust. This process was pushing the two sides of the ridge apart and this is the

mechanism that had separated, and indeed is still separating, the African and South

American continents, which was first understood by the German scientist Alfred

Wegener but failed to provide the physical explanation.

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Figure 1.7. Structure of the mid-oceanic ridges in offset segments and with

a deep central gorge (Menard, 1976).

The new oceanic crust that is being continually formed at the mid-oceanic ridges

should expand the Earth unless there is another mechanism that consumes this newly

formed material. Seismologists found that in certain parts of the world earthquakes

occur at large depths below the lithosphere. The lithosphere is the only part of the

planet that is brittle and can fracture. Therefore these deep events suggest that

somehow in such regions the lithosphere is descending into the mantle, being

consumed at the same rate that new crust is being generated at the oceanic ridges.

This process that occurs where two plates collide and one is pushed down below the

other, is known as subduction.

The evidence provided by the structure of the mid-ocean ridges and the location of

seismic activity at both the mid-ocean ridges and in subduction zones was used to

formulate the theory of plate tectonics (e.g., Isacks et al., 1968; McKenzie, 1968). As

indicated previously, the Earth’s surface is divided into a number of lithospheric slabs

called tectonic plates and they move relative to each other as a result of the underlying

convection currents in the mantle (Figure 1.8). The plates interact at their boundaries

in one of three ways (Figure 1.9). At the ocean ridges, plates move apart from eachother and they are divergent plate boundaries. At convergent boundaries (where two

plates collide), one plate will usually be driven below the other in the process of

subduction. Oceanic plate is subducted below continental plate along the Pacific coast

of South and Central America; oceanic crust is subducted below oceanic crust in the

Caribbean arc. When two continental plates collide, there is enormous deformation

and thickening of the lithosphere along the boundary (e.g., the Himalayas). Two

plates can also move horizontally, pass one another at transcurrent boundaries. Such

boundaries are can be seen along long and well-defined faults such as the San

Andreas Fault in California, which is the boundary between the North American and

Pacific plates, and the North Anatolian Fault in Turkey. Figure 1.10 shows the global

distribution of the three different types of plate boundaries that can be related to themotions depicted in Figure 1.9.

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Figure 1.8. Directions of motion of the principal tectonic plates (Dowrick, 1987).

Figure 1.9. Schematic illustration of the mechanisms of interaction between tectonic plates

(Shearer, 1999).

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Figure 1.10. The main lithospheric plates and the nature of their boundaries (Udías, 1999).

The majority of earthquakes and volcanoes around the world can be linked to the

interaction of tectonic plates. Figure 1.11 shows sea-floor spreading and subduction;

as the oceanic plate descends in the subduction trench it is heated and eventually

reaches its fusion point. The lower density material of the oceanic crust with respect

to its surrounding material melts and rises towards the surface emerging as lava toform volcanoes. This is the mechanism of most of the volcanoes and that is why the

volcanoes are generally found above the subducted plate.

Figure 1.11. Illustration of subduction causing volcanism (Press & Siever, 1986).

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Figure 1.11 also illustrates that most earthquakes are located at the boundaries of

tectonic plates. These earthquakes are formed with the forces generated by the

interaction of the tectonic plates. This fact explains why the dense bands of seismicity

around the world coincide with the boundaries between tectonic plates (Figure 1.1)

Earthquake activity zones at depth in the subducted slab are generally referred to as

Wadati-Benioff zones and they can occur at depths of as much as 700 km. Theintraplate earthquakes, such as those in the New Madrid zone (Figure 1.2), that occur

far from the tectonic plate boundaries have more complex mechanisms. It is unlikely

that the triggering factor for these events is the increased crustal stresses due to plate

interactions as is the case at the plate margins. A number of different explanations

have been put forward as possible reasons for the generation of earthquakes in stable

continental regions: stress concentrations around zones of weakness, release of crustal

stresses due to deglaciation or reduction of the mechanical strength of crustal rocks

due to the action of fluids or of heat (Johnston et al., 1994).

1.2 Formation of earthquakes

The dynamics of the Earth’s lithosphere and plate tectonics presented in the previous

section explain how the relative movements of the tectonic slabs create deformation,

and earthquake activity along the major boundaries. The actual mechanism by which

earthquakes occur can be explained by the elastic rebound theory that is introduced

after the 1906 San Francisco earthquake by Reid (1911).

The US earth scientists studied the 1906 San Francisco earthquake in great detail

(Lawson, 1908). The fault rupture that was traced for a distance of more than 400 km

along the line of the San Andreas showed a predominant horizontal slip with features

traversed by the fault trace, such as fences (Figure 1.12) and roads (Figure 1.13). The

slip between the two sides of the fault line varied between 2 m to 4 m on average.

Figure 1.12. Lateral offset of a fence caused by rupture on the San Andreas fault in the 1906

San Francisco earthquake (Lawson, 1908).

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Figure 1.13. Offset of a road caused by the fault rupture (Lawson, 1908).

The investigation of the measured displacements along the fault line as well as the re-

triangulation (theodolite angle measurements) of the survey points in northern

California and data from earlier triangulations in 1851-1865 and again in 1874-1892

were re-examined during these studies. It was found that during the decades prior to

the earthquake, points on opposite sides of the fault had moved relative to one

another. Both the directions and the magnitudes of these displacements were

consistent with the observed slip along the fault line after the earthquake.

On the basis of these observations, the leader of the group charged with investigating

the fault rupture, Harry Fielding Reid, proposed the theory of elastic rebound as the

mechanism for earthquake occurrence, a model that has been universally accepted.

Figure 1.14 illustrates the elastic rebound theory. As plates on opposite sides of a fault

are subjected to force and shift, they accumulate energy and slowly deform until their

internal strength is exceeded. At that time, a sudden movement occurs along the fault,

releasing the accumulated energy, and the rocks snap back to their original

undeformed shape. The elastic rebound theory was the first theory to satisfactorily

explain earthquakes. Previously it was thought that ruptures of the surface were the

result of strong ground shaking rather than the converse suggested by this theory.

With the exception of volcanic earthquakes that are the result of sudden and massive

movements of magma (and which also tend to be limited to fairly small events), all

earthquakes are caused by rupture on geological faults. The rupture begins at one

particular point and then propagates along the fault plane very rapidly: average

velocities of rupture are between 2 and 3 km/s.

http://en.wikipedia.org/wiki/Theory

http://en.wikipedia.org/wiki/Theory

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Figure 1.14. Schematic illustration of the mechanism of elastic rebound theory:

the block represents a portion of the crust within which there is a geological fault (the thin

black line). Due to the movement and interaction of tectonic plates the crust deforms, storing

elastic energy while the fault is effectively ‘locked’. When the resistance of the fault isexceeded, it ruptures allowing the two sides to move abruptly back towards their original

positions, relaxing the crust in the process.

The fault ruptures are often very complex but they can be idealized as rectangular

planes. The crustal block above the fault plane is called as the hanging wall, the block

below the fault plane is the footwall (Figure 1.15). The angle of the maximum slope

of the fault plane, measured downwards from the horizontal in the vertical plane is the

dip angle, ° (Figure 1.16). The strike, °, is the angle between line of intersection

of the fault plane and the ground surface. It is measured clockwise from north in the

horizontal plane. The conventional way to measure the strike angle is that the hanging

wall is on the right (Figure 1.16).

Ruptures involving only horizontal motion are called strike-slip faults. They are

classified as either right-lateral (dextral) or left-lateral (sinistral) as illustrated inFigure 1.15. If the movement on the fault is purely vertical the rupture is referred to as

dip-slip, which may be either normal (hanging wall moves downwards) or reverse

(hanging wall moves upwards). Normal faults are the result of tectonic extension

(when two slabs diverge). The reverse faults are encountered in zones of compression

(when two slabs converge) such as the Caucasus (Figure 3.17). Reverse faults with

shallow angles of dip (less than about 30°) are known as thrust faults. Many faults

show a combination of both horizontal and vertical slip and these are known as

oblique ruptures (Figure 1.15).

The direction of slip on a rupture can be defined by the angle known as the rake or

slip, °. The rake is the angle between the strike direction and the vector of slip of

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the hanging wall with respect to the fault wall. It is measured in the plane of the fault,

positive upwards. A normal fault rupture has a rake of –90° and a reverse rupture a

rake of +90°.

Figure 1.15. Mechanisms of fault rupture (Reiter, 1990).

Figure 1.16. Illustration of the definition of the dip ( ), strike ( ) and rake ( )

of a fault rupture (Shearer, 1999).

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Figure 1.17. View of the zone of maximum slip on the causative fault of the 1988 Spitak

earthquake in Armenia (Bommer & Ambraseys, 1989). Notice the three

figures to the left of the picture for scale.

The basic measures of the size of an earthquake are magnitude and seismic moment

that are introduced in the next section. The size of the fault rupture area increases as

the size of an earthquake increases. This is shown in Figure 1.18. For earthquakes of magnitude 6, Figure 1.18 suggests that the average area of the fault rupture will be of

the order of 75 km2

that would correspond to circle of about 5 km radius. The ruptures

of events of this size and smaller may be approximately circular and such events can

be idealized as point sources, but for magnitudes above 6 the rupture will tend to

become rectangular and these events are represented as extended sources.

Rectangular fault ruptures can be characterised by their length, L, and their width, W.

The length is measured along the strike and the width is measured along down the dip

of the fault plane. Many studies have produced empirical regressions between

dimensions of the fault rupture, or the slip, and the earthquake magnitude by

performing regressions on the observed rupture dimensions, slip and magnitude(Figure 1.19). The paper by Wells and Coppersmith (1994) is the most referred one

among these studies. The relationships of Wells & Coppersmith (1994) are provided

in Tables 1.1 and 1.2. The equations should only be used to obtain estimates of the

value on the left-hand side from known values on the right-hand side of the equation.

Empirical correlations between the dimensions of fault ruptures, their slip and

earthquake magnitude are very useful for performing seismic hazard assessment.

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Figure 1.18. Correlation between the area of the fault rupture and earthquake magnitude

(Reiter, 1990).

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Figure 1.19. Regression relationships between earthquake magnitude and maximum surface

displacement (Reiter, 1990) based on the worldwide data set of Slemmons (1982).

Table 1.1. Regressions of rupture length, rupture width, rupture area, and moment

magnitude (Mw) from Wells & Coppersmith (1994).

Table 1.2. Regressions of displacement and moment magnitude (Mw)

from Wells & Coppersmith (1994).

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1.3 Seismic waves

An earthquake can be defined as a sudden release of elastic strain energy in the

Earth’s crust as the result of fault rupture. This energy radiates from the fault rupture

and from the surrounding volume of the crust in the form of seismic waves. The

seismic waves within an elastic medium can propagate by one of two mechanisms:

compression or shear. The two basic types of seismic waves corresponding to these

two mechanisms are P-waves (P stands for primary because they travel faster arrive

first to any point of observation) and S-waves (secondary). P-waves are

compressional waves that carry energy either by compression or dilation. They are

classified as longitudinal waves because the particle motion that carries the energy

forward takes place in the direction of wave propagation. S-waves are shear waves

and they are transverse because the particle motion is perpendicular to the direction of

wave propagation. S-waves are often considered resolved into two components, SH

and SV, which are polarised in the horizontal and vertical planes respectively. The

mechanisms of propagation of P- and S-waves, collectively known as body waves, are

illustrated in Figure 1.20.

The propagation velocity of P- and S-waves, Vp and Vs respectively, depends only onthe elastic properties of the medium in which they travel:

)21)(1(

)1( E V p

(1.1)

G

)1(2

E V p

(1.2)

where E is the modulus of elasticity, is the mass density, is Poisson’s ratio, and

G is the shear modulus. The values of the wave velocities generally increase with

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depth within the Earth with values within the crust typically of the order of 6 km/s for

Vp and 4 km/s for Vs. P-waves usually travel about 3 times faster than S-waves.

Figure 1.20. Mechanism of propagation of body waves.

As some of the body waves reach the Earth’s surface and are reflected back

downwards, they cause perturbations of the surface that give rise to a second type of wave (surface waves) that propagate along the surface of the Earth. The two basic

types of surface waves are Love waves (LQ) and Rayleigh waves (LR). The

amplitude of these waves is a maximum at the surface, rapidly decreasing with depth.

Unlike body waves, whose propagation velocity depends only on the elastic

properties of the medium, surface wave velocities additionally depend on their

periods. The mechanisms of surface wave propagation are illustrated in Figure 1.21.

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Figure 1.21. Mechanism of propagation of surface waves.

1.4. Magnitude of an earthquake

Magnitude is the most familiar measurement of earthquake size. It measures the

total amount of seismic energy released in an earthquake. The magnitudeconcept was first introduced by Richter (1935) to quantify the size of

earthquakes in southern California. Richter measured the maximum amplitude

A (mm) of recordings on Wood-Anderson seismographs and defined the

magnitude of the earthquake as:

(1.3)

where A0 is the amplitude that an earthquake of magnitude zero (arbitrarily

defined size of the earthquake that would produce a maximum trace amplitude

of 0.001 mm on a Wood-Anderson instrument at a distance of 100 km). Richter

presented tables of -log(A0) as a function of epicentral distances up to 1,000 km.

Magnitudes calculated in this way are referred to as local magnitudes and

represented by the symbol ML. Figure 1.22 shows a chart that can be employed

to determine the value of ML graphically from readings of amplitude and S- and

P-wave arrival interval on a Wood-Anderson seismograph. The chart presents

average conditions in southern California for ML calculation.

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Figure 1.22. Nonogram for the rapid estimation of local magnitude (Willmore, 1979).

Most national and regional seismograph networks calculate local magnitudes, someusing the original ML definition of Richter (1935), others deriving their own scales

that reflect the regional attenuation characteristics and the use of seismographs other

than the Wood-Anderson model.

The concept of local magnitude presented by Richter (1935) is limited in

application because it is calibrated specifically for southern California or for any

other particular region where it is applied. Moreover the original definition of

ML is based specifically on the Wood-Anderson seismograph and is limited to

earthquakes recorded at distances of less than 1,000 km. All of these

disadvantages are overcome by the use of teleseismic magnitude scales that are

based not on the amplitude of the recorded trace but rather on the maximumratio of the amplitude A (measured in m, i.e. 10-6 m) to period T (seconds) of

the seismic instrument and is therefore independent of the type of seismograph.

In order to obtain the appropriate value of A it is necessary to know the

magnification factor of the instrument at the period T.

There are two teleseismic magnitude scales defined, one based on

measurements of body waves, mb, and another based on measurements of

surface-waves, Ms. The equation for calculating surface-wave magnitude is:

(1.4)

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The equation for calculating body-wave magnitude is:

(1.5)

Table 1.3. Attenuation function, Q, for shallow earthquake (Willmore, 1979).

For shallow earthquakes, the value of Q is read from tables such as Table 1.3. For

deeper earthquakes the appropriate value of Q is read from curves that present Q as a

function of both epicentral distance ° and focal depth h, an example of which is

shown in Figure 1.23. Unless correction is made for depth, Ms values for deep

earthquakes (which produce weak surface waves) will be much lower than mb.

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Figure 1.23. Values of Q as function of epicentral distance and focal depth for the vertical

component of the direct P phase (Willmore, 1979).

Ms and mb are calculated from long-period and short-period seismographs,respectively and therefore they correspond to the amplitude of the seismic waves in

different parts of the spectrum. Thus, local and teleseismic magnitude scales do not

necessarily agree. Empirical relationships exist that allow values on one scale to be

estimated from values of another. A typical set of equations for such an objective is

derived by Ambraseys and Bommer (1990) that are given below. The standard

deviations associated with these regressions are significant and hence these

conversions should be used with caution.

51.0 M 66 .0m75.0 L b 2.0 (1.6)

91.1 M 50.0m87 .0 sb 22.0 (1.7)

20.1 M 58 .0 M 82.0 s L 21.0 (1.8)

All of the magnitude scales described so far have a fundamental shortcoming: they are

unable to distinguish the relative size of the largest earthquakes, a phenomenon

known as saturation. Hanks and Kanamori (1979) defined another magnitude scale

that is called as moment magnitude (represented by either M or Mw). Moment

magnitude does not suffer from saturation because it is related to the seismic moment

(M0) that directly yields the total amount of energy released by the earthquake

rupture. Eq. (1.9) gives the moment magnitude vs. seismic moment relationshipprovided by Hanks and Kanamori (1979).

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0.6 ) M log(3

2 M ow (5.14)

Figure 1.24 compares the moment magnitude scale with the teleseismic and local

magnitude scales, illustrating the phenomenon of magnitude saturation. The mB scaleis similar to mb but calculated at slightly longer periods; MJMA is the magnitude scale

of the Japanese Meteorological Society, which is calculated from the ground-motion

amplitudes measured from medium period seismographs.

Figure 1.24. Comparison of moment magnitude scale with other magnitude scales (Reiter,

1990).

1.5. Intensity of an earthquake

Damage on structures due to strong ground shaking can be measured by observations

or accelerograms recorded by accelerographs. Accelerographs record the acceleration

of particles as a function of time. Unlike many other seismic instruments they only

record strong shaking and they must be installed in the vicinity of active seismic

sources. The accelerographs generally record three mutually perpendicular

components of motion in the vertical and two horizontal directions (Figure 1.25).

Accelerograms contain significant information about the nature of the ground shaking

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and also about the highly varied characteristics that different earthquakes can produce

at different locations (Figure 1.26). Ground-motion parameters (e.g., peak ground

acceleration, velocity or spectral ordinates) that are obtained from the accelerograms

quantitatively describe the state of damage and are used for the seismic design and

performance assessment of structural systems. In addition to the information about the

variation of acceleration with time during the earthquake, double integration of theaccelerogram (through some special processing) provides the velocity and

displacement time-histories as well (Figure 1.27).

Figure 1.25. Typical triaxial strong-motion record with acceleration time-histories in two

horizontal directions and the vertical

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Figure 1.26. Accelerograms from various earthquakes and recording sites, illustrating the

variability in the nature of strong ground-motion (Hudson, 1979).

Figure 1.27. Accelerogram from the 1978 Tabas earthquake in Iran, with velocity and

displacement time-hsitories (Bommer & Elnashai, 1999).

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The first accelerographs are analog instruments that record on film or paper. The most

popular of these instruments is the SMA-1 (Figures 1.28 and 1.29). There are a

number of shortcomings with first generation accelerographs. Firstly, they trigger

after a certain threshold value and they generally cannot capture the very first wave

arrivals. Although these are generally not of much interest to engineers, it does mean

that the initial values of velocity and displacement (obtained by integrating theacceleration data) are unknown. Secondly, the recorded waveforms are poor in

resolution. Thirdly, the recorded waveforms should be digitized for their advance

usage and digitization always introduces noise to the original waveform.

Figure 1.28. Photograph of an SMA-1 optical-mechanical accelerograph.

Figure 1.29. Typical analog record from an SMA-1 accelerograph obtained on 70 mm film

(Hudson, 1979).

The second-generation accelerographs operate with a force-balance transducer and

record digitally on to solid state or magnetic media. These instruments can record the

waveforms in higher resolution due to their wide dynamic range and there is no need

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to digitize the record. Furthermore, they can operate continuously and the first

motions of the earthquake shaking are retained (Figure 1.30).

Figure 1.30. Recordings of an earthquake in Italy on optical mechanical (top) and digital

(bottom) accelerographs at the same site.

The qualitative measurement of strong ground-shaking influence on structural

systems as well as on the earthquake area is done through macroseismic intensity; an

index that reflects the strength of ground shaking at a particular location during an

earthquake. This definition clearly indicates that the macroseismic intensity is not

really a measure of the size of the earthquake in the same way as magnitude. The

macroseismic intensity is usually represented by Roman numerals. Thus, it is not acontinuous variable and can never attain decimal values. Intensity VIII indicates a

level of shaking that may cause damage in engineered structures; stronger shaking

that does not quite qualify as intensity IX may legitimately be represented as VIII-IX,

VIII+ or IX-, but to write 8.5 is incorrect.

Macroseismic intensity scales were first devised at the beginning of the 19th

century.

Rossi and Forel formally presented the first intensity scale in 1883. Many of the

earlier scales have now been abandoned and the main scales remained are those of the

Japanese Meteorological Agency (JMA), used in Japan, the Modified Mercalli (MM)

scale, used in the Americas, and the Medvedev-Karnik-Sponheuer (MSK) scale used

in most of Europe, except Italy where the Mercalli-Cancani-Sieberg scale (MCS) isemployed. Figure 1.31 shows a comparison of these scales.

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Figure 1.31. Comparison of intensity scales (Reiter, 1990).

The MSK scale is now superseded by the European Macroseismic Scale (Grunthal,

1998), known by its acronym EMS. It is an attempt to correct some of the

inconsistencies and shortcomings in the earlier scale. An abbreviated form of the

EMS is shown in Table 1.4. The MM and MSK scales, as well as others, base the

grades of intensity on (a) the shaking felt by humans, (b) the movement of objects, (c)

the damage to buildings, and (d) changes in the ground. This last category is not a

reliable indicator because of the very high heterogeneity of soils and rocks. It is

generally advised not to use observations of ground behaviour other than for

corroboration with the MM and MSK scales; the EMS actually removes geotechnical

and geological observations from the descriptions of different grades of intensity.

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Table 1.4. Summary version of European Macroseismic Scale (Grunthal, 1998)

The lower levels of intensity are defined primarily by how people feel the shaking but

as values increase human perception of the movement becomes progressively less

important as damage to buildings becomes more prominent. It is worth keeping inmind some key points along the scales that help to make it much easier to interpret

and to assess intensities. Firstly, intensity III is generally considered as the threshold

of perceptibility, below which the ground shaking is not felt by most people. Intensity

VII can be thought of as the threshold of appreciable building damage, although this

might be intensity VIII for engineered structures. Intensity XII is very rarely, if ever,

encountered in reality and therefore XI can be treated as the upper bound. In practice,

X appears to be an effective upper bound.

One final point must be made: the scales are neither continuous nor linear. The

variation from degree to degree is not gradual, each increase in one degree

representing a jump in the level of shaking. Furthermore, the jumps between different

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degrees are not equal: the increase in the level of ground shaking from IV to V is not

the same and the increase from VII to VIII.

Macroseismic intensity is evaluated by direct field observations. Macroseismic

intensity is sometimes thought of as an indication of building damage, but this is

really a misperception. Different types of buildings are treated differently and levelsof damage are carefully defined precisely to deconvolute the ground motion from the

response of buildings. Table 1.5 shows how EMS classifies the vulnerability

(susceptibility of structures to damage) of different types of building. Figures 1.32

and 1.33 illustrate the different grades of damage. Table 1.5 presents the detailed

descriptions of the degrees of intensity VII, VIII and IX on the EMS scale. These are

the grades that describe the most important degrees of shaking in many ways, from an

engineering point, since they are frequently encountered and potentially damaging.

They each include (a) descriptions of people’s reactions, (b) the movement of objects

and (c) the proportions of each building category expected to be within each damage

state.

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Table 1.5. Definition of structural vulnerability classes in EMS (Grunthal, 1998).

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Figure 1.32. Classification of damage to masonry buildings in EMS (Grunthal,1998).

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Figure 1.33. Classification of damage to reinforced concrete buildings in EMS(Grunthal,1998).

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Table 1.6. Descriptions of intensity degrees in EMS (Grunthal, 1998).

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