CE232 Earthquake Engineering Report

8/10/2019 CE232 Earthquake Engineering Report

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CE 232 Earthquake Engineering

SEISMOLOGY AND EARTHQUAKE REPORT

Submitted by

Eduardo M. Oroz Jr.


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TABLE OF CONTENTS

Internal Structure of the Earth.3

Continental Drift and Plate Tectonics..5

Rupture of fault and Earthquakes.7

Other sources of Earthquakes..10

Location of Earthquakes.12

Size and Measurement of Earthquakes13

Modification of Earthquake due to the Nature of Soil..16

Earthquake Damage to Structures..17


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Internal Structure of the Earth

Three Parts of Earth's Interior

Knowledge of earth's interior is essential for understanding plate tectonics. A good

analogy for teaching about earth's interior is a piece of fruit with a large pit such as a peach or aplum. Most students are familiar with these fruits and have seen them cut in half.

If we cut a piece of fruit in half we will see that it is composed of three parts: 1) a very

thin skin, 2) a seed of significant size located in the center, and 3) most of the mass of the fruit

being contained within the flesh. Cutting the earth we would see: 1) a very thin crust on the

outside, 2) a core of significant size in the center, and 3) most of the mass of the Earth contained

in the mantle.

Earth's Crust

There are two different types of crust: thin oceanic crust that underlies the ocean basins

and thicker continental crust that underlies the continents. These two different types of crust are

made up of different types of rock. The thin oceanic crust is composed of primarily of basalt and

the thicker continental crust is composed primarily of granite. The low density of the thick

continental crust allows it to "float" in high relief on the much higher density mantle below.


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Earth's Mantle

Earth's mantle is thought to be composed mainly of olivine-rich rock. It has different

temperatures at different depths. The temperature is lowest immediately beneath the crust and

increases with depth. The highest temperatures occur where the mantle material is in contact

with the heat-producing core. This steady increase of temperature with depth is known as thegeothermal gradient. The geothermal gradient is responsible for different rock behaviors and the

different rock behaviors are used to divide the mantle into two different zones. Rocks in the

upper mantle are cool and brittle, while rocks in the lower mantle are hot and soft (but not

molten). Rocks in the upper mantle are brittle enough to break under stress and produce

earthquakes. However, rocks in the lower mantle are soft and flow when subjected to forces

instead of breaking. The lower limit of brittle behavior is the boundary between the upper and

lower mantle.

Earth's Core

Earth's Core is thought to be composed mainly of an iron and nickel alloy. This

composition is assumed based upon calculations of its density and upon the fact that many

meteorites are iron-nickel alloys. The core is earth's source of internal heat because it contains

radioactive materials which release heat as they break down into more stable substances.

The core is divided into two different zones. The outer core is a liquid because the

temperatures there are adequate to melt the iron-nickel alloy. However, the inner core is a solid

even though its temperature is higher than the outer core. Here, tremendous pressure, produced

by the weight of the overlying rocks is strong enough to crowd the atoms tightly together and

prevents the liquid state.


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Continental Drift and Plate Tectonics

Continental Drift

Continental drift is the movement of the Earth's continents relative to each other, thus

appearing to drift across the ocean bed. The speculation that continents might have 'drifted' was

first put forward by Abraham Ortelius in 1596. The concept was independently and more fully

developed by Alfred Wegener in 1912, but his theory was rejected by some for lack of a

mechanism (though this was supplied later by Holmes) and others because of prior theoretical

commitments. The idea of continental drift has been subsumed by the theory of plate tectonics,

which explains how the continents move.

Plate Tectonics

The continents drift slowly (the timescale for substantial change is 10-100 million years),

but that they drift at all is remarkable. The following figure illustrates the structure of the first

100-200 kilometers of the Earth's interior, and provides an answer to this question.

The crust is thin, varying from a few tens of kilometers thick beneath the continents to to

less than 10 km thick beneath the many of the oceans. The crust and upper mantle together

constitute the lithosphere, which is typically 50-100 km thick and is broken into large plates (not

illustrated). These plates sit on the aesthenosphere.

The aesthenosphere is kept plastic (deformable) largely through heat generated by

radioactive decay. The material that is decaying is primarily radioactive isotopes of light

elements like aluminum and magnesium. This heat source is small on an absolute scale (the

corresponding heat flow at the surface out of the Earth is only about 1/6000 of the solar energy


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falling on the surface). Nevertheless, because of the insulating properties of the Earth's rocks this

is sufficient to keep the aesthenosphere plastic in consistency.

Plate tectonics (from the Late Latin tectonicus, from the Greek: "pertaining to

building") is a scientific theory that describes the large-scale motion of Earth's lithosphere. This

theoretical model builds on the concept of continental drift which was developed during the firstfew decades of the 20th century. The geoscientific community accepted the theory after the

concepts of seafloor spreading were later developed in the late 1950s and early 1960s.

The lithosphere and the aesthenosphere

The lithosphere, which is the rigid outermost shell of a planet (on Earth, the crust and

upper mantle), is broken up into tectonic plates. On Earth, there are seven or eight major plates

(depending on how they are defined) and many minor plates. Where plates meet, their relative

motion determines the type of boundary; convergent, divergent, or transform. Earthquakes,

volcanic activity, mountain-building, and oceanic trench formation occur along these plate

boundaries. The lateral relative movement of the plates typically varies from zero to 100 mm

annually.

Tectonic plates are composed of oceanic lithosphere and thicker continental lithosphere,

each topped by its own kind of crust. Along convergent boundaries, subduction carries plates

into the mantle; the material lost is roughly balanced by the formation of new (oceanic) crust

along divergent margins by seafloor spreading. In this way, the total surface of the globe remains

the same. This prediction of plate tectonics is also referred to as the conveyor belt principle.


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Earlier theories (that still have some supporters) propose gradual shrinking (contraction) or

gradual expansion of the globe.

Tectonic plates are able to move because the Earth's lithosphere has greater strength than

the underlying asthenosphere. Lateral density variations in the mantle result in convection. Plate

movement is thought to be driven by a combination of the motion of the seafloor away from thespreading ridge (due to variations in topography and density of the crust, which result in

differences in gravitational forces) and drag, with downward suction, at the subduction zones.

Another explanation lies in the different forces generated by the rotation of the globe and the

tidal forces of the Sun and Moon. The relative importance of each of these factors and their

relationship to each other is unclear, and still the subject of much debate.

Rupture of Fault and Earthquakes

Rupture of Faults

A fault is a break in the earth's crust along which movement can take place causing an

earthquake. In Utah, movement along faults is mostly vertical; mountain blocks (for example, the

Wasatch Range) move up relative to the downward movement of valley blocks (for example, the

Salt Lake Valley).

Faults with evidence of Holocene (about 10,000 years ago to present) movement are the

main concern because they are most likely to generate future earthquakes. If the earthquake is

large enough, surface fault rupture can occur.

With a large earthquake (about magnitude 6.5 and greater), the fault rupture can reach

and displace the ground surface, forming a fault scarp (steep break in slope). The resulting faultscarp may be several inches to 20 feet in height, and up to about 40 miles in length, depending

on the size of the earthquake.

An area hundreds of feet wide can be affected, called the zone of deformation, which

occurs chiefly on the downthrown side of the main fault and encompasses multiple minor faults,

cracks, local tilting, and grabens (downdropped blocks between faults). Buildings in the zone of

deformation would be damaged, particularly those straddling the main fault.

Also, anything crossing the fault, such as transportation corridors, utilities, and other

lifelines, both underground and above ground, can be damaged or broken. The ground can be

dropped below the water table on the downdropped side, resulting in localized flooding.

Surface fault rupture can also cause tectonic subsidence, which is the broad, permanent

tilting of the valley floor down toward the fault scarp. Tilting can cause flooding along lake and

reservoir shorelines nearest the fault; along altered stream courses; and along canals, sewer lines,

or other gravity-flow systems where slope gradients are lessened or reversed.


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Earthquake fault types

There are three main types of fault, all of which may cause an interplate earthquake:

normal, reverse (thrust) and strike-slip. Normal and reverse faulting are examples of dip-slip,

where the displacement along the fault is in the direction of dip and movement on them involves

a vertical component. Normal faults occur mainly in areas where the crust is being extended suchas a divergent boundary. Reverse faults occur in areas where the crust is being shortened such as

at a convergent boundary. Strike-slip faults are steep structures where the two sides of the fault

slip horizontally past each other; transform boundaries are a particular type of strike-slip fault.

Many earthquakes are caused by movement on faults that have components of both dip-slip and

strike-slip; this is known as oblique slip.

Reverse faults, particularly those along convergent plate boundaries are associated with

the most powerful earthquakes, megathrust earthquakes, including almost all of those of

magnitude 8 or more. Strike-slip faults, particularly continental transforms can produce major

earthquakes up to about magnitude 8. Earthquakes associated with normal faults are generallyless than magnitude 7.

This is so because the energy released in an earthquake, and thus its magnitude, is

proportional to the area of the fault that ruptures and the stress drop. Therefore, the longer the

length and the wider the width of the faulted area, the larger the resulting magnitude. The

topmost, brittle part of the Earth's crust, and the cool slabs of the tectonic plates that are

descending down into the hot mantle, are the only parts of our planet which can store elastic

energy and release it in fault ruptures. Rocks hotter than about 300 degrees Celsius flow in

response to stress; they do not rupture in earthquakes. The maximum observed lengths of

ruptures and mapped faults, which may break in one go are approximately 1000 km. Examplesare the earthquakes in Chile, 1960; Alaska, 1957; Sumatra, 2004, all in subduction zones. The

longest earthquake ruptures on strike-slip faults, like the San Andreas Fault (1857, 1906), the

North Anatolian Fault in Turkey (1939) and the Denali Fault in Alaska (2002), are about half to

one third as long as the lengths along subducting plate margins, and those along normal faults are

even shorter.

The most important parameter controlling the maximum earthquake magnitude on a fault

is however not the maximum available length, but the available width because the latter varies by

a factor of 20. Along converging plate margins, the dip angle of the rupture plane is very

shallow, typically about 10 degrees. Thus the width of the plane within the top brittle crust of theEarth can become 50 to 100 km (Japan, 2011; Alaska, 1964), making the most powerful

earthquakes possible.

Strike-slip faults tend to be oriented near vertically, resulting in an approximate width of

10 km within the brittle crust, thus earthquakes with magnitudes much larger than 8 are not

possible. Maximum magnitudes along many normal faults are even more limited because many


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of them are located along spreading centers, as in Iceland, where the thickness of the brittle layer

is only about 6 km.

In addition, there exists a hierarchy of stress level in the three fault types. Thrust faults

are generated by the highest, strike slip by intermediate and normal faults by the lowest stress

levels. This can easily be understood by considering the direction of the greatest principal stress,the direction of the force that 'pushes' the rock mass during the faulting. In the case of normal

faults, the rock mass is pushed down in a vertical direction, thus the pushing force (greatest

principal stress) equals the weight of the rock mass itself. In the case of thrusting, the rock mass

'escapes' in the direction of the least principal stress, namely upward, lifting the rock mass up,

thus the overburden equals the least principal stress. Strike-slip faulting is intermediate between

the other two types described above. This difference in stress regime in the three faulting

environments can contribute to differences in stress drop during faulting, which contributes to

differences in the radiated energy, regardless of fault dimensions.

Earthquake

An earthquake (also known as a quake, tremor or temblor) is the result of a sudden

release of energy in the Earth's crust that creates seismic waves. The seismicity, seismism or

seismic activity of an area refers to the frequency, type and size of earthquakes experienced over

a period of time.

Earthquakes are measured using observations from seismometers. The moment

magnitude is the most common scale on which earthquakes larger than approximately 5 are

reported for the entire globe. The more numerous earthquakes smaller than magnitude 5 reported

by national seismological observatories are measured mostly on the local magnitude scale, alsoreferred to as the Richter magnitude scale. These two scales are numerically similar over their

range of validity. Magnitude 3 or lower earthquakes are mostly almost imperceptible or weak

and magnitudes 7 and over potentially cause serious damage over larger areas, depending on

their depth. The largest earthquakes in historic times have been of magnitude slightly over 9,

although there is no limit to the possible magnitude. The most recent large earthquake of

magnitude 9.0 or larger was a 9.0 magnitude earthquake in Japan in 2011 (as of March 2014),

and it was the largest Japanese earthquake since records began. Intensity of shaking is measured

on the modified Mercalli scale. The shallower an earthquake, the more damage to structures it

causes, all else being equal.

At the Earth's surface, earthquakes manifest themselves by shaking and sometimes

displacement of the ground. When the epicenter of a large earthquake is located offshore, the

seabed may be displaced sufficiently to cause a tsunami. Earthquakes can also trigger landslides,

and occasionally volcanic activity.

In its most general sense, the word earthquake is used to describe any seismic event

whether natural or caused by humans that generates seismic waves. Earthquakes are caused


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mostly by rupture of geological faults, but also by other events such as volcanic activity,

landslides, mine blasts, and nuclear tests. An earthquake's point of initial rupture is called its

focus or hypocenter. The epicenter is the point at ground level directly above the hypocenter.

Other Sources of Earthquakes

Naturally occurring earthquakes

Tectonic earthquakes occur anywhere in the earth where there is sufficient stored elastic

strain energy to drive fracture propagation along a fault plane. The sides of a fault move past

each other smoothly and aseismically only if there are no irregularities or asperities along the

fault surface that increase the frictional resistance. Most fault surfaces do have such asperities

and this leads to a form of stick-slip behaviour. Once the fault has locked, continued relative

motion between the plates leads to increasing stress and therefore, stored strain energy in the

volume around the fault surface. This continues until the stress has risen sufficiently to break

through the asperity, suddenly allowing sliding over the locked portion of the fault, releasing thestored energy. This energy is released as a combination of radiated elastic strain seismic waves,

frictional heating of the fault surface, and cracking of the rock, thus causing an earthquake. This

process of gradual build-up of strain and stress punctuated by occasional sudden earthquake

failure is referred to as the elastic-rebound theory. It is estimated that only 10 percent or less of

an earthquake's total energy is radiated as seismic energy. Most of the earthquake's energy is

used to power the earthquake fracture growth or is converted into heat generated by friction.

Therefore, earthquakes lower the Earth's available elastic potential energy and raise its

temperature, though these changes are negligible compared to the conductive and convective

flow of heat out from the Earth's deep interior.

Earthquakes away from plate boundaries

Where plate boundaries occur within the continental lithosphere, deformation is spread

out over a much larger area than the plate boundary itself. In the case of the San Andreas fault

continental transform, many earthquakes occur away from the plate boundary and are related to

strains developed within the broader zone of deformation caused by major irregularities in the

fault trace (e.g., the "Big bend" region). The Northridge earthquake was associated with

movement on a blind thrust within such a zone. Another example is the strongly oblique

convergent plate boundary between the Arabian and Eurasian plates where it runs through the

northwestern part of the Zagros mountains. The deformation associated with this plate boundaryis partitioned into nearly pure thrust sense movements perpendicular to the boundary over a wide

zone to the southwest and nearly pure strike-slip motion along the Main Recent Fault close to the

actual plate boundary itself. This is demonstrated by earthquake focal mechanisms.

All tectonic plates have internal stress fields caused by their interactions with

neighbouring plates and sedimentary loading or unloading (e.g. deglaciation). These stresses may

be sufficient to cause failure along existing fault planes, giving rise to intraplate earthquakes.


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Shallow-focus and deep-focus earthquakes

The majority of tectonic earthquakes originate at the ring of fire in depths not exceeding

tens of kilometers. Earthquakes occurring at a depth of less than 70 km are classified as 'shallow-

focus' earthquakes, while those with a focal-depth between 70 and 300 km are commonly termed

'mid-focus' or 'intermediate-depth' earthquakes. In subduction zones, where older and colderoceanic crust descends beneath another tectonic plate, deep-focus earthquakes may occur at

much greater depths (ranging from 300 up to 700 kilometers). These seismically active areas of

subduction are known as Wadati-Benioff zones. Deep-focus earthquakes occur at a depth where

the subducted lithosphere should no longer be brittle, due to the high temperature and pressure.

A possible mechanism for the generation of deep-focus earthquakes is faulting caused by olivine

undergoing a phase transition into a spinel structure.

Rupture dynamics

A tectonic earthquake begins by an initial rupture at a point on the fault surface, a processknown as nucleation. The scale of the nucleation zone is uncertain, with some evidence, such as

the rupture dimensions of the smallest earthquakes, suggesting that it is smaller than 100 m while

other evidence, such as a slow component revealed by low-frequency spectra of some

earthquakes, suggest that it is larger. The possibility that the nucleation involves some sort of

preparation process is supported by the observation that about 40% of earthquakes are preceded

by foreshocks. Once the rupture has initiated it begins to propagate along the fault surface. The

mechanics of this process are poorly understood, partly because it is difficult to recreate the high

sliding velocities in a laboratory. Also the effects of strong ground motion make it very difficult

to record information close to a nucleation zone.

Rupture propagation is generally modeled using a fracture mechanics approach, likening

the rupture to a propagating mixed mode shear crack. The rupture velocity is a function of the

fracture energy in the volume around the crack tip, increasing with decreasing fracture energy.

The velocity of rupture propagation is orders of magnitude faster than the displacement velocity

across the fault. Earthquake ruptures typically propagate at velocities that are in the range 70

90% of the S-wave velocity and this is independent of earthquake size. A small subset of

earthquake ruptures appear to have propagated at speeds greater than the S-wave velocity. These

supershear earthquakes have all been observed during large strike-slip events. The unusually

wide zone of coseismic damage caused by the 2001 Kunlun earthquake has been attributed to the

effects of the sonic boom developed in such earthquakes. Some earthquake ruptures travel atunusually low velocities and are referred to as slow earthquakes. A particularly dangerous form

of slow earthquake is the tsunami earthquake, observed where the relatively low felt intensities,

caused by the slow propagation speed of some great earthquakes, fail to alert the population of

the neighbouring coast, as in the 1896 Meiji-Sanriku earthquake.


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Tidal forces

Research work has shown a robust correlation between small tidally induced forces and non-

volcanic tremor activity.

Earthquake clusters

Most earthquakes form part of a sequence, related to each other in terms of location and

time. Most earthquake clusters consist of small tremors that cause little to no damage, but there is

a theory that earthquakes can recur in a regular pattern.

Aftershocks

An aftershock is an earthquake that occurs after a previous earthquake, the mainshock.

An aftershock is in the same region of the main shock but always of a smaller magnitude. If an

aftershock is larger than the main shock, the aftershock is redesignated as the main shock and the

original main shock is redesignated as a foreshock. Aftershocks are formed as the crust aroundthe displaced fault plane adjusts to the effects of the main shock.

Earthquake swarms

Earthquake swarms are sequences of earthquakes striking in a specific area within a short

period of time. They are different from earthquakes followed by a series of aftershocks by the

fact that no single earthquake in the sequence is obviously the main shock, therefore none have

notable higher magnitudes than the other. An example of an earthquake swarm is the 2004

activity at Yellowstone National Park. In August 2012, a swarm of earthquakes shook Southern

California's Imperial Valley, showing the most recorded activity in the area since the 1970s.

Earthquake storms

Sometimes a series of earthquakes occur in a sort of earthquake storm, where the

earthquakes strike a fault in clusters, each triggered by the shaking or stress redistribution of the

previous earthquakes. Similar to aftershocks but on adjacent segments of fault, these storms

occur over the course of years, and with some of the later earthquakes as damaging as the early

ones. Such a pattern was observed in the sequence of about a dozen earthquakes that struck the

North Anatolian Fault in Turkey in the 20th century and has been inferred for older anomalous

clusters of large earthquakes in the Middle East.

Location of Earthquakes

The primary purpose of a seismometer is to locate the initiating points of earthquake

epicenters. The secondary purpose, of determining the 'size' or Moment magnitude scale must be

calculated after the precise location is known.


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The earliest seismographs were designed to give a sense of the direction of the first

motions from an earthquake. The Chinese frog seismograph would have dropped its ball in the

general compass direction of the earthquake, assuming a strong positive pulse. We now know

that first motions can be in almost any direction depending on the type of initiating rupture (focal

mechanism).

The first refinement that allowed a more precise determination of the location, was the

use of a time scale. Instead of merely noting, or recording, the absolute motions of a pendulum,

the displacements were plotted on a moving graph, driven by a clock mechanism. This was the

first seismogram, which allowed precise timing of the first ground motion, and an accurate plot

of subsequent motions.

From the first seismograms, as seen on the figure, it was noticed that the trace was

divided into two major portions. The first seismic wave to arrive was the P-wave, followed

closely by the S-wave. Knowing the relative 'velocities of propagation', it was a simple matter to

calculate the distance of the earthquake.

One seismograph would give the distance, but that could be plotted as a circle, with an

infinite number of possibilities. Two seismographs would give two intersecting circles, with two

possible locations. Only with a third seismograph would there be a precise location.

The process of accurate location, was greatly improved with the advent of precise

absolute timing. Early seismographs were almost always located at an astronomical observatory,

just for the purpose of timing. See the history of the Canadian Dominion Observatory, is also the

Geological Survey of Canada seismology laboratory. Recently, GPS is being used for accurate

time, and seismometers can be located almost anywhere.

Modern earthquake location still requires a minimum of three seismometers. Most likely,

there are many, forming a seismic array. The emphasis is on precision, since much can be

learned about the fault mechanics and seismic hazard, if the locations can be determined to

within a kilometer or two, for small earthquakes. For this, computer programs use an iterative

process, involving a 'guess and correction' algorithm. As well, a very good model of the local

crustal velocity structure is required: seismic velocities vary with the local geology. For P-waves,

the relation between velocity and bulk density of the medium has been quantified in Gardner's

relation.

Size and Measurement of Earthquakes

Measuring and locating earthquakes

Earthquakes can be recorded by seismometers up to great distances, because seismic

waves travel through the whole Earth's interior. The absolute magnitude of a quake is

conventionally reported by numbers on the moment magnitude scale (formerly Richter scale,


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magnitude 7 causing serious damage over large areas), whereas the felt magnitude is reported

using the modified Mercalli intensity scale (intensity IIXII).

Every tremor produces different types of seismic waves, which travel through rock with

different velocities:

Longitudinal P-waves (shock- or pressure waves)

Transverse S-waves (both body waves)

Surface waves(Rayleigh and Love waves)

Propagation velocity of the seismic waves ranges from approx. 3 km/s up to 13 km/s,

depending on the density and elasticity of the medium. In the Earth's interior the shock- or P

waves travel much faster than the S waves (approx. relation 1.7 : 1). The differences in travel

time from the epicentre to the observatory are a measure of the distance and can be used to image

both sources of quakes and structures within the Earth. Also the depth of the hypocenter can becomputed roughly.

In solid rock P-waves travel at about 6 to 7 km per second; the velocity increases within

the deep mantle to ~13 km/s. The velocity of S-waves ranges from 23 km/s in light sediments

and 45 km/s in the Earth's crust up to 7 km/s in the deep mantle. As a consequence, the first

waves of a distant earthquake arrive at an observatory via the Earth's mantle.

On average, the kilometer distance to the earthquake is the number of seconds between

the P and S wave times 8.[46] Slight deviations are caused by inhomogeneities of subsurface

structure. By such analyses of seismograms the Earth's core was located in 1913 by Beno

Gutenberg.

Earthquakes are not only categorized by their magnitude but also by the place where they

occur. The world is divided into 754 FlinnEngdahl regions (F-E regions), which are based on

political and geographical boundaries as well as seismic activity. More active zones are divided

into smaller F-E regions whereas less active zones belong to larger F-E regions.

Standard reporting of earthquakes includes its magnitude, date and time of occurrence,

geographic coordinates of its epicenter, depth of the epicenter, geographical region, distances to

population centers, location uncertainty, a number of parameters that are included in USGS

earthquake reports (number of stations reporting, number of observations, etc.), and a uniqueevent ID.

Size of Earthquakes

The magnitude is the most often cited measure of an earthquake's size, but it is not the

only measure, and in fact, there are different types of earthquake magnitude. Early estimates of

earthquake size were based on non-instrumental measures of the earthquakes effects. For


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example, we could use values such as the number of fatalities or injuries, the maximum value of

shaking intensity, or the area of intense shaking. The problem with these measures is that they

don't correlate well. The damage and devastation produced by an earthquake will depend on its

location, depth, proximity to populated regions, as well as its "true" size. Even for earthquakes

close enough to population centers values such as maximum intensity and the area experiencing

a particular level of shaking did not correlate well.

In 1931 a Japanese seismologist named Kiyoo Wadati constructed a chart of maximumground motion versus distance for a number of earthquakes and noted that the plots for differentearthquakes formed parallel, curved lines (the larger earthquakes produced larger amplitudes).

The fact that earthquakes of different size generated curves that were roughly parallel suggested

that a single number could quantify the relative size of different earthquakes.

In 1935 Charles Richter constructed a similar diagram of peak ground motion versus

distance and used it to create the first earthquake magnitude scale (a logarithmic relationship

between earthquake size and observed peak ground motion). He based his scale on an analogywith the stellar brightness scale commonly used in astronomy which is also similar to the pH

scale used to measure acidity (pH is a logarithmic measure of the Hydrogen ion concentration in

a solution).

To complete the construction of the magnitude scale, Richter had to establish a reference

value and identify the rate at which the peak amplitudes decrease with distance from an

earthquake. He established a reference value for earthquake magnitude when he defined the

magnitude as the base-ten logarithm of the maximum ground motion (in micrometers) recorded

on a Wood-Anderson short-period seismometer one hundred kilometers from the earthquake.

Richter was pragmatic in his definition, and chose a value for a magnitude zero that insured that

most of the earthquakes routinely recorded would have positive magnitudes. Also, the Wood-

Anderson short-period instrument that Richter chose for his reference records seismic waves

with a period of about 0.8 seconds, roughly the vibration periods that we feel and that damage

our buildings and other structures.

Richter also developed a distance correction to account for the variation in maximum

ground motion with distance from an earthquake (the dashed curves shown in the above diagram

show his relationship for southern California). The precise rate that the peak ground motions

decrease with distance depends on the regional geology and thus the magnitude scale for

different regions is slightly dependent on the "distance correction curve".

Thus originally, Richter's scale was specifically designed for application in southern

California. Richter's method became widely used because it was simple, required only the

location of the earthquake (to get the distance) and a quick measure of the peak ground motion,

was more reliable than older measures such as intensity. It became widely used, well established,

and forms the basis for many of the measures that we continue to use today. Generally the


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magnitude is computed from seismographs from as many seismic recording stations as are

available and the average value is used as our estimate of an earthquake's size.

We call the Richter's original magnitude scale ML (for "local magnitude"), but the press

usually reports all magnitudes as Richter magnitudes.

Modification of Earthquake due to Nature of Soil

Soil Conditions can have considerable effect on the shape of the surface motion response

during earthquakes. Since the changes in the shape of the response spectra can have significant

effects on the lateral forces on structures, it would seem desirable that the nature of soil

conditions underlying a site be taken into account when evaluating the lateral forces for design

purposes. It is also apparent that the most accurate information available on the dynamic

properties of such soils should be used. While the information produced by Seed and Idriss has

been shown to differ from that determined for this study, the caution that their data is only

intended to be used as a guideline and considerable judgments may have to be exercised. Theysuggest using a range of material properties in the computations when equipment for the accurate

determination of moduli and equivalent viscous damping factors is not available.

The various site properties are extremely involved in controlling the response, since with

increasing input acceleration the amplification factor decreases due to the higher damping

employed in the softer soils. The highest amplification factors are generally associated with the

weaker motions. This phenomenon has been observed by many researchers and writers.

One aspect which cannot be bypassed is the influence of the characteristics of the base

rock motion on ground response. Insufficient evidence is available to state categorically whether

bedrock motions should be characterized by response spectra having a flat or a peakedform.

Many writers believe the former to be a better representation. Regardless of the general form of

the spectra the fact should not be overlooked that seemingly small peaks in the bedrock spectra

are able to excite higher modes of vibration of the soil deposit producing a multi-peaked surface

response. This phenomenon often occurs and in some cases the peak spectral value may be

developed at a period corresponding to the predominant period of the bedrock motion. Perhaps

the answer is to use a range of possible but realistic input motions to bracket the range of input

accelerations.

The non-linear soil properties give rise to a surface motion response spectrum whose

peak spectral value and predominant period is dependent oon the magnitude of the maximum

base rock acceleration. In a truly linear system, scaling the maximum input acceleration merely

scales the peak spectral response by the same factor. The non-linearity of the soil strength

properties precludes this simple scaling as a marked change in predominant period of the surface

motion results and an analysis must be made to determine the full effect. In other words there is

no simple relationship relating maximum input acceleration to predominant period and peak

spectral velocity. Quantitatively, it is generally apparent, that deposits of deep soft soils tend to


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produce ground motions having long period characteristics, while shallower deposits of stiff soils

tend to produce shorter period ground motions.

Clearly the level of understanding of surface layer modification of earthquake shear

waves has reached the point where engineers may now use the results of research as a useful

guide to evaluate the influence of soil condition on earthquake ground motions.

Earthquake Damage to Structures

Comparatively speaking, the absolute movement of the ground and buildings during an

earthquake is not actually all that large, even during a major earthquake. That is, they do not

usually undergo displacements that are large relative to the building's own dimensions. So, it is

not the distance that a building moves which alone causes damage.

Rather, it is because a building is suddenly forced to move very quickly that it suffers

damage during an earthquake. Think of someone pulling a rug from beneath you. If they pull it

quickly (i.e., accelerate it a great deal), then they needn't pull it very far to throw you off balance.

On the other hand, if they pull the rug slowly and only gradually increase the speed of the rug,

they can move (displace) it a great distance without that same unfortunate result.

In other words, the damage that a building suffers primarily depends not upon its

displacement, but upon acceleration. Whereas displacement is the actual distance the ground and

the building may move during an earthquake, acceleration is a measure of how quickly they

change speed as they move. During an earthquake, the speed at which both the ground and

building are moving will reach some maximum. The more quickly they reach this maximum, the

greater their acceleration.

It's worthwhile mentioning here that in order to study the earthquake responses ofbuildings; many buildings in earthquake-prone regions of the world have been equipped withstrong motion accelerometers. These are special instruments which are capable of recording the

accelerations of either the ground or building, depending upon their placement.

The recording of the motion itself is known as an accelerogram. Figure below shows an

accelerogram recorded in a hospital building parking lot during the Northridge, Californiaearthquake of January 17, 1994.


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In addition to providing valuable information about the characteristics of the particular

earthquake recorded or the building where the accelerogram was recorded, accelerograms

recorded in the past are also often used in the earthquake response analysis and earthquakedesign of buildings yet to be constructed.

Newton's Law

Acceleration has this important influence on damage, because, as an object in movement,

the building obeys Newton' famous Second Law of Dynamics. The simplest form of the equationwhich expresses the Second Law of Motion is F = MA.

This states the Force acting on the building is equal to the Mass of the building times the

Acceleration. So, as the acceleration of the ground, and in turn, of the building, increase, so does

the force which affects the building, since the mass of the building doesn't change.

Of course, the greater the force affecting a building, the more damage it will suffer;

decreasing F is an important goal of earthquake resistant design. When designing a new building,for example, it is desirable to make it as light as possible, which means, of course, that M, and inturn, F will be lessened. As we've seen in the discussion of Advanced Earthquake Resistant

Techniques, various techniques are now also available for reducing A.

Inertial Forces

Acceleration, Inertial Forces

It is important to note that F is actually what's known as an inertial force, that is, the force

is created by the building's tendency to remain at rest, and in its original position, even thoughthe ground beneath it is moving. This is in accordance with another important physical law

known as D'Alembert's Principle, which states that a mass acted upon by acceleration tends tooppose that acceleration in an opposite direction and proportionally to the magnitude of the

acceleration.

This inertial force F imposes strains upon the building's structural elements. These

structural elements primarily include the building's beams, columns, load-bearing walls, floors,as well as the connecting elements that tie these various structural elements together. If these

strains are large enough, the building's structural elements suffer damage of various kinds.


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Building Frequency and Period

To begin with, as we discussed in the How Earthquakes Affect Buildings, the magnitudeof the building responsethat is, the accelerations which it undergoes depends primarily upon

the frequencies of the input ground motion and the building's natural frequency. When these are

near or equal to one another, the building's response reaches a peak level.

In some circumstances, this dynamic amplification effect can increase the building

acceleration to a value two times or more that of the ground acceleration at the base of thebuilding. Generally, buildings with higher natural frequencies, and a short natural period, tend to

suffer higher accelerations but smaller displacement. In the case of buildings with lower natural

frequencies, and a long natural period, this is reversed as the buildings will experience loweraccelerations but larger displacements.

Building Stiffness

The taller a building, the longer its natural period tends to be. But the height of a buildingis also related to another important structural characteristic: the building flexibility. Tallerbuildings tend to be more flexible than short buildings. (Only consider a thin metal rod. If it is

very short, it is difficulty to bend it in your hand.

If the rod is somewhat longer, and of the same diameter, it will become much easier to

bend. Buildings behave similarly.) We say that a short building is stiff, while a taller building is

flexible. (Obviously, flexibility and stiffness are really just the two sides of the same coin. Ifsomething is stiff, it isn't flexible and vice-versa.)

Stiffness greatly affects the building's uptake of earthquake generated force. Reconsider

our first example above, of the rigid stone block deeply founded in the soil. The rigid block ofstone is very stiff; as a result it responds in a simple, dramatic manner. Real buildings, of course,

are more inherently flexible, being composed of many different parts.

Furthermore, not only is the block stiff, it is brittle; and because of this, it cracks during

the earthquake. This leads us to the next important structural characteristic affecting a building'searthquake response and performance and ductility.

Ductility

Ductility is the ability to undergo distortion or deformation bending, for example

without resulting in complete breakage or failure. To take once again the example of the rigidblock in Figure 3, the block is an example of a structure with extremely low ductility. To see

how ductility can improve a building's performance during an earthquake, consider figure below.


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Metal Rod Ductility

For the block, we have substituted a combination of a metal rod and a weight. In response

to the ground motion, the rod bends but does not break. (Of course, metals in general are more

ductile than materials such as stone, brick and concrete.) Obviously, it is far more desirable for a

building to sustain a limited amount of deformation than for it to suffer a complete breakage

failure.

The ductility of a structure is in fact one of the most important factors affecting its

earthquake performance. One of the primary tasks of an engineer designing a building to be

earthquake resistant is to ensure that the building will possess enough ductility to withstand the

size and types of earthquakes it is likely to experience during its lifetime.

Damping

The last of the important structural characteristics, or parameters, which we'll discuss

here is damping. As we noted earlier, ground and building motion during an earthquake has a

complex, vibratory nature. Rather than undergoing a single "yank" in one direction, the building

actually moves back and forth in many different horizontal directions

All vibrating objects, including buildings, tend to eventually stop vibrating as time goes

on. More precisely, the amplitude of vibration decays with time. Without damping, a vibrating

object would never stop vibrating, once it had been set in motion. Obviously, different objects

possess differing degrees of damping. A bean bag, for example, has high damping; a trampoline

has low damping.

In a building undergoing an earthquake, damping the decay of the amplitude of a

building's vibrations is due to internal friction and the absorption of energy by the building's

structural and nonstructural elements. All buildings possess some intrinsic damping.

The more damping a building possesses, the sooner it will stop vibrating--which of

course is highly desirable from the standpoint of earthquake performance. Today, some of the

more advanced techniques of earthquake resistant design and construction employ added

damping devices like shock absorbers to increase artificially the intrinsic damping of a building

and so improve its earthquake performance.


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References

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

http://eqseis.geosc.psu.edu/~cammon/HTML/Classes/IntroQuakes/Notes/earthquake_size.html

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

http://www.isr.umd.edu/~austin/aladdin.d/matrix-eq-spectra.html

http://geology.utah.gov/utahgeo/hazards/eqfault/eqfault.htm

http://mceer.buffalo.edu/infoservice/reference_services/buildingRespondEQ.asp

Documents

CE232 Earthquake Engineering Report