seismic and the earth

8/13/2019 Seismic and the Earth

1/32

http://www.as.uky.edu/sites/default/files/elearning/module06

swf.swf

http://www.as.uky.edu/sites/default/files/elearning/module06

swf.swf

Seismic and the Earth's Structure

The structure of Earth's deep interior cannot be studied directly. But geologists use seismic

(earthquake) waves to determine the depths of layers of molten andsemi-molten material within Earth. Because different types of earthquake waves behave

differently when they encounter material in different states (for example,molten, semi-molten, solid), seismic stations established around Earth detect and record thestrengths of the different types of waves and the directions from which

they came. Geologists use these records to establish the structure of Earth's interior.

The two principal types of seismic waves are P-waves (pressure; goes through liquid and solid)

and S-waves (shear or secondary; goes only through solid - not through liquid). The travel

velocity of these two wave types is not the same (P-waves are faster than S-waves). Thus, ifthere is an earthquake somewhere, the first waves that arrive are P-waves. In essence, the gap in

P-wave and S-wave arrival gives a first estimate of the distance to the earthquake.


2/32

Above image shows some typical seismograms with arrival of P- and S-waves marked.

As we know from physics, all waves change direction when they pass through layers of differentdensity (refraction). That is what makes light collect in a magnifying glass, and that is also what

makes seismic waves travel in curved paths through the Earth (because of the increasing

pressure, materials are more dense towards the core, travel velocity of seismic waves increases).Refraction of seismic waves causes them to curve away from a direct path. Reflection causes

them to glance off certain surfaces (e.g. core mantle boundary) when they hit it at too shallow of

an angle. The result of this behavior, in combination with the fact that S-waves can not travelthrough liquids, is the appearance of seismic shadows, opposite of the actual earthquake site.

The geometric distribution and extent of these shadows as measured for a given earthquake(many receiver stations - seismographs, are needed all over the world to do that) allows us to

calculate the position of major boundaries in the Earth's interior, as well as giving us informationabout the solid vs liquid character of the various layers, and even about some of their physicalproperties.


3/32

The biggest discontinuity in the Earth's interior is the core/mantle boundary, because there wehave a strong density contrast between the iron core (density between 10-11 g/cm

3) and the

silicate mantle (density from 3.3-5.5 g/cm3, increases with depth).

Background sound is the actual recording of an earthquake.

Seismology and Earth's Interior

There are two categories of earthquake waves. Body wavescan travel deep into theEarth; Surface wavescan only travel very near the surface of the Earth. There are twokinds of body waves, and two kinds of surface waves. As you might imagine, onlybody waves can give us any information about the deep interior of the Earth.

All earthquakes are relatively shallow, with the deepest at about 700 km depth. Anearthquake generates body waves that spread out in all directions, like light from anaked light bulb. Notice in the diagram below that you can think of earthquake waves asmoving out like rays(arrows) or as wave fronts(spherical shells). Surface wave raystravel out in all horizontal directions (like the arrows on the top of the block picturedbelow), like ripples moving out from a pebble dropped into a pond.


4/32

All over the surface of the Earth are seismograph stations which can detect all of thewaves that arrive at that location. By recognizing what kinds of waves have arrived,exactly when they arrived, and knowing where and when the earthquake occurred (orsometimes the earthquake location and time itself is determined by seismographstations), we can learn about the deep interior of the Earth. This is because thesewaves refract (bend) and reflect at boundaries in the Earth.

Strain is dimensionless; Stress has units of force/unit area. 1

Newton/m2= 1 Pascal = 1 Pa

Body Waves:

There are two kinds of body waves corresponding to the two fundamental ways you candeform an object: you can squeeze it (or stretch it, which is like "negative squeezing"),or you can shear it.

P Waves

The diagram on the left above illustrates aP wave.These are also calledcompressionalor longitudinalwaves. Material is compressed and stretched in the
http://principles.ou.edu/glossary.html#p%20wavehttp://principles.ou.edu/glossary.html#p%20wavehttp://principles.ou.edu/glossary.html#p%20wavehttp://principles.ou.edu/glossary.html#p%20wave


5/32

horizontal direction, from left to right, and the wave (disturbance) also travels in thehorizontal direction. P waves travel faster than any other type of wave. They can travelthrough fluid or solid materials. Ordinary sound waves in air are P waves.

P comes from primarywave, because they arrive first, but a mnemonic is push-pull wave

P wave velocity depends on a material's "plane wave modulus" and its density:

Where is Lam's constant, is shear modulus, K is bulk modulus, and is density. Notice that

density is in the denominator, so denser rocks should be slower. However, although the density

of rock in the Earth generally increases with depth, the rigidity, as expressed in the variouselastic constants, increases even more rapidly with depth. Hence, P wave velocity generally

increases with increasing depth.

Since solids, liquids and gasses have a finite bulk modulus, P waves can travel through any of

these

S Waves

The diagram on the right above illustrates anS wave.These are also called shearwaves. S comes from secondary wave. Material is sheared, so that an imaginary

square drawn on the side of the block becomes diamond shaped. The material vibratesup and down (or side to side, in and out of the screen, if the hammer had struck the sideof the block instead of the top) but the wave (disturbance) travels in the horizontaldirection from left to right. S waves travel more slowly than P waves. They can onlytravel through solid materials. Plucking a guitar string generates a kind of shearwave; the string vibrates side to side, but the wave travels along the string.

S-wave velocity depends on a material's shear modulus, , and density, :

Since fluids (liquids and gasses have zero shear modulus, S waves cannot travel throughfluids.However, seismic waves have a period no larger than minutes. Some materials, like the

mantle, are solids on that time scale, but not on the time scale of millions of years.

Comparing the velocity expressions, you can see thatVP> VSfor any material.
http://principles.ou.edu/glossary.html#s%20wavehttp://principles.ou.edu/glossary.html#s%20wavehttp://principles.ou.edu/glossary.html#s%20wavehttp://principles.ou.edu/glossary.html#s%20wave


6/32

For both types of body waves:

P and S waves travel faster in rigid, dense rocks. Rocks generally get more rigidand denser with depth. Generally, though, elastic constants increase morerapidly than density, so the velocity of P and S waves generally increases

with depth. P and S waves are refracted and reflected at boundaries. In the diagram below, the subsurface earthquake location (focus, or hypocenter

) is shown in yellow. The ray we've shown coming out of the earthquake travelsin a straight line in the blue layer. When it reaches the red layer (which might beslower or faster), the ray splits: some of the energy goes into the red layer but isbent (refracted), and some of the energy is reflected back up to the surface. Ananalogy: When you stand in front of a store window, you can usually see yourreflection, proving that some of the light reflects back at you. But people in thestore can also see you, so some of the light goes through the glass. [Reflectivitycan be calculated from Zoeppritz equations.]

Surface Waves

waves which travel only along the Earth's surface amplitude decreases exponentially with depth (relative to wavelength)

thus, short wavelength have shallow penetration similar to skin effect of electromagnetic waves traveling along conductive medium

(submarine comms) Historically, recognized as "large" waves, hence designated by "L"

Love Waves

LQ(Quer: German for lateral) horizontally polarized shear waves predicted by A.E.H. Love in 1911


7/32


8/32

Sign determines whether polarity reversal occurs:

In upper crust, changes in sometimes small, the refl ection coeff icientoften depends mainlyonvelocity dif ferences. (Just a rule of thumb.)

Refractions

Refractions occur when velocities differ (if they don't, ray passes through unbent!):

Snell's Law


9/32

Snell's law applies to reflectionsand refractions, even with mode conversion:


10/32

In large regions of the Earth, velocity increase gradually with depth, leading to gradualbending of rays; where there are abrupt velocity changes, sharp bending, and reflections,

will occur.


11/32

These reflected and refracted rays show up as different phaseson a seismogram. Here is a


12/32

simple one:

Earthquake Seismology and the Interior of

the Eartth


13/32

The main points about using earthquakes waves to determine the internal structure ofthe Earth are summarized here, then explained in more detail:

By measuring travel times of earthquake waves to seismograph stations, we candetermine velocity structure of Earth

By making graphs of travel time versus distance between earthquakes andseismograph stations, we findo velocity generally increases gradually w/ depth in Earth, due to increasing

pressure and rigidity of the rockso however, there are abrupt velocity changes at certain depths, indicating

layering The 4 major layers in the Earth, from outside in, are the crust, mantle, outer core,

and inner core.o Thecrustis very thin, averaging about 30 km thick in the continents and 5

km thick in the oceanso Themantleis 2900 km thick (almost halfway to the center of the Earth. It

is made of dark, dense, ultramafic rock (peridotite).o Theouter coreis 2300 km thick and is made of a mixture liquid iron

(90%) and nickel (10%)o Theinner coreis at the center of the Earth and has a 1200 km radius; it's

made of solid iron (90%) and nickel (10%).

Crust - Mantle Boundary

The crust mantle boundary was discovered in 1909 by a seismologist namedMohorovici (Yugoslav), as a result of his study of an earthquake in Croatia at thattime.

He found that, out to about 150 km, the time it took for the earthquake waves toreach each seismograph station was proportional to the distance the station wasfrom the earthquake. He used the familiar time/distance/rate equation (distance =rate*time, or rate = distance/time) to determine that the velocity of the upper crustmust be about 6 km/s. In the graph below, this corresponds to the straight linesegment on the left, which has a slope of corresponding to 6 km/s.

However, for stations greater than about 150 km from the earthquake, waves didnot take as much longer to arrive as if they were traveling at only 6 km/s. In fact,the slope of the second line segment corresponds to a velocity of 8 km/s.
http://principles.ou.edu/glossary.html#crusthttp://principles.ou.edu/glossary.html#crusthttp://principles.ou.edu/glossary.html#crusthttp://principles.ou.edu/glossary.html#mantlehttp://principles.ou.edu/glossary.html#mantlehttp://principles.ou.edu/glossary.html#mantlehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#mantlehttp://principles.ou.edu/glossary.html#crust


14/32

Furthermore, Mohorovici figured out that the distance at which the change inslope occurred (about 150 km) can be used to calculate the depth to velocityincrease from 6 to 8 km/s. He calculated that the depth to this velocity jump wasabout 30 km.

We interpret this velocity jump as the crust-mantle boundary, and often refer to itas theMohorovicic discontinuity,or Moho, for short.

The diagram below shows a cross-section of the crust and mantle, with theearthquake on the left. The triangles on the surface are meant to be seismographstations at different distances from the earthquake. At short distances, the "directwaves" that travel along the surface will arrive first. However, at greaterdistances, the waves that travel down to the mantle, and are bent and travelalong the top of the mantle at the higher velocity, can arrive before the wavestraveling directly along the surface. These refracted waves make up for the extradistance by traveling faster for most of their path.

Seismic refraction experiments like Mohorovici's have been, and still are, beingconducted all over the Earth. They indicate that continental crust is about 35 km
http://principles.ou.edu/glossary.html#Mohorovicichttp://principles.ou.edu/glossary.html#Mohorovicichttp://principles.ou.edu/glossary.html#Mohorovicichttp://principles.ou.edu/glossary.html#Mohorovicic


15/32

thick, but varies greatly from place to place, and oceanic crust is pretty uniformly5 km thick.

This contour map of the thickness of the Earth's crust was developed from the CRUST 5.1

model. The contour interval is 10 km; we also include the 45 km contour for greater detail on thecontinents.

Jeffreys-Bullen travel-time curve

Another global raypath diagram

Seismogram, near Fiji, GOL, 92.2 deg, 6.4 Mb

Core - Mantle Boundary

The core-mantle boundary was discovered in 1913 by a seismologist namedGutenberg. Seismologists had noticed that P waves are not recorded atseismograph stations which are from 104

oto 140

oaway from an earthquake (the
http://principles.ou.edu/eq_seismo/jbcurve.gifhttp://principles.ou.edu/eq_seismo/jbcurve.gifhttp://principles.ou.edu/eq_seismo/raypaths3.jpghttp://principles.ou.edu/eq_seismo/raypaths3.jpghttp://principles.ou.edu/eq_seismo/seismogram_fiji_s.jpghttp://principles.ou.edu/eq_seismo/seismogram_fiji_s.jpghttp://principles.ou.edu/eq_seismo/seismogram_fiji_s.jpghttp://principles.ou.edu/eq_seismo/seismogram_fiji_s.jpghttp://principles.ou.edu/eq_seismo/raypaths3.jpghttp://principles.ou.edu/eq_seismo/jbcurve.gif


16/32

angle is the angle made by drawing a line from the earthquake to the center ofthe Earth, and then from there to the seismograph station.

Gutenberg explained this Shadow Zone with acorewhich slowed and bent Pwaves

Later, anS wave shadow zonewas recognized, meaning no S waves werereceived at seismographs stations from 104

oto 180

ofrom an earthquake; the S

wave shadow zone is caused by the outer core, which is liquid iron/nickel.
http://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/glossary.html#corehttp://principles.ou.edu/shadow_zone.gifhttp://principles.ou.edu/shadow_zone.gifhttp://principles.ou.edu/shadow_zone.gifhttp://principles.ou.edu/shadow_zone.gifhttp://principles.ou.edu/glossary.html#core


17/32

Modeling of seismic waves traveling through the Earth allowed seismologists todetermine that the core begins at a depth of 2900 km, or in other words, themantleis 2900 km thick; its composition is probablyultramaficrock (peridotite).This is based on the velocity of the waves, meteorites, mass of the Earth andother lines of evidence.

Inner Core - Outer Core Boundary

In 1936, a Swedish seismologist named Inge Lehmann recognized waves whichwere reflected from a boundary deep within the Earth. She correctly interpretedthis as the outside of the inner core, which is solid iron and nickel.

In the 1960's, nuclear blasts allowed for a more precise determination of theradius of the innner core. U.S.'s nuclear blasts were always at a known spot, andwere detonated exactly at a specified time. This eliminated much of theuncertainty seismologists have to deal with with natural earthquakes, whoseprecise origin time and location must be worked out by the travel timesthemselves!

Copyri ght 2009 J. L. Ahern
http://principles.ou.edu/glossary.html#mantlehttp://principles.ou.edu/glossary.html#mantlehttp://principles.ou.edu/glossary.html#ultramafichttp://principles.ou.edu/glossary.html#ultramafichttp://principles.ou.edu/glossary.html#ultramafichttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#peridotitehttp://principles.ou.edu/glossary.html#ultramafichttp://principles.ou.edu/glossary.html#mantle


18/32

Prof. Stephen A. Nelson EENS 111

Tulane University Physical Geology

Earthquakes and the Earth's Interior

This page last updated on 24-Oct-2003

Earthquakes

Earthquakes occur when energy stored in elastically strained rocks is suddenly released. Thisrelease of energy causes intense ground shaking in the area near the source of the earthquake

and sends waves of elastic energy, called seismic waves, throughout the Earth. Earthquakes canbe generated by bomb blasts, volcanic eruptions, and sudden slippage along faults. Earthquakesare definitely a geologic hazard for those living in earthquake prone areas, but the seismic

waves generated by earthquakes are invaluable for studying the interior of the Earth.

Origin of Earthquakes

Most natural earthquakes are caused by suddenslippage along a fault zone. The elastic

rebound theorysuggests that if slippage alonga fault is hindered such that elastic strain

energy builds up in the deforming rocks on

either side of the fault, when the slippage doesoccur, the energy released causes an

earthquake. This theory was discovered by

making measurements at a number of pointsacross a fault. Prior to an earthquake it was

noted that the rocks adjacent to the fault were

bending. These bends disappeared after an

earthquake suggesting that the energy stored in

bending the rocks was suddenly releasedduring the earthquake.

Seismology, The Study of Earthquakes


19/32

When an earthquake occurs, the elastic energy is released and sends out vibrations that travelthroughout the Earth. These vibrations are called seismic waves. The study of how seismicwaves behave in the Earth is called seismology.

Seismographs - Seismic

waves travel through the

Earth as vibrations. Aseismometeris an

instrument used to

record these vibrations

and the resulting graphthat shows the

vibrations is called a

seismograph. Theseismometer must be

able to move with the

vibrations, yet part of itmust remain nearly

stationary.

This is accomplished by isolating the recording device (like a pen) from the rest of the

Earth using the principal of inertia. For example, if the pen is attached to a large mass

suspended by a spring, the spring and the large mass move less than the paper which isattached to the Earth, and on which the record of the vibrations is made.

Seismic Waves. The source of an

earthquake is called the focus,

which is an exact location withinthe Earth were seismic waves are

generated by sudden release of

stored elastic energy. The epicenteris the point on the surface of the

Earth directly above the focus.

Sometimes the media get these two

terms confused. Seismic wavesemanating from the focus can travel

in several ways, and thus there are

several different kinds of seismicwaves.


20/32

o Body Waves - emanate

from the focus and

travel in all directionsthrough the body of the

Earth. There are two

types of body waves:

P - waves- are Primary waves. They travel with a velocity that dependson the elastic properties of the rock through which they travel.

Vp= [(K + 4/3)/]Where, Vpis the velocity of the P-wave, K is the incompressibility of the

material, is the rigidity of the material, and is the density of thematerial.

P-waves are the same thing as sound waves. They move through the

material by compressing it, but after it has been compressed it expands,so that the wave moves by compressing and expanding the material as it

travels. Thus the velocity of the P-wave depends on how easily the

material can be compressed (the incompressibility), how rigid thematerial is (the rigidity), and the density of the material. P-waves have

the highest velocity of all seismic waves and thus will reach all

seismographs first.

S-Waves- Secondary waves, also called shear waves. They travel with a

velocity that depends only on the rigidity and density of the material

through which they travel:

Vs= [( )/]S-waves travel through material by shearing it or changing its shape in

the direction perpendicular to the direction of travel. The resistance to

shearing of a material is the property called the rigidity. It is notable thatliquids have no rigidity, so that the velocity of an S-wave is zero in a

liquid. (This point will become important later). Note that S-waves travel


21/32

slower than P-waves, so they will reach a seismograph after the P-wave.

o Surface Waves- Surface waves differ from body waves in that they do not

travel through the Earth, but instead travel along paths nearly parallel to the

surface of the Earth. Surface waves behave like S-waves in that they cause up

and down and side to side movement as they pass, but they travel slower than S-waves and do not travel through the body of the Earth.

The record of an

earthquake, aseismograph, as

recorded by a

seismometer, will

be a plot ofvibrations versus

time. On the

seismograph, time ismarked at regular

intervals, so that we

can determine the

time of arrival ofthe first P-wave and

the time of arrival

of the first S-wave.

(Note again, that because P-waves have a higher velocity than S-waves, the P-waves arrive at

the seismographic station before the S-waves).


22/32

Location of Earthquakes - Inorder to determine the location

of an earthquake, we need to

have recorded a seismographof the earthquake from at least

three seismographic stations at

different distances from theepicenter of the quake. In

addition, we need one further

piece of information - that isthe time it takes for P-waves

and S-waves to travel through

the Earth and arrive at a

seismographic station. Such

information has been collectedover the last 80 or so years,

and is available as travel time

curves.

From the seismographs at eachstation one determines the S-P

interval (the difference in the

time of arrival of the first S-wave and the time of arrival of

the first P-wave. Note that on

the travel time curves, the S-P

interval increases withincreasing distance from the

epicenter. Thus the S-P interval

tells us the distance to the

epicenter from theseismographic station where

the earthquake was recorded.

Thus, at each station we candraw a circle on a map that has

a radius equal to the distance


23/32

from the epicenter.

Three such circles will intersect in a point that locates the epicenter of the earthquake.

Magnitude of Earthquakes - Whenever a large destructive earthquake occurs in theworld the press immediately wants to know where the earthquake occurred and how big

the earthquake was (in California the question is usually - Was this the Big One?). Thesize of an earthquake is usually given in terms of a scale called the Richter Magnitude.

Richter Magnitude is a scale of earthquake size developed by a seismologist named

Charles F. Richter. The Richter Magnitude involves measuring the amplitude (height)of the largest recorded wave at a specific distance from the earthquake. While it is

correct to say that for each increase in 1 in the Richter Magnitude, there is a tenfold

increase in amplitude of the wave, it is incorrectto say that each increase of 1 in

Richter Magnitude represents a tenfold increase in the size of the Earthquake (as iscommonly incorrectly stated by the Press).

A better measure of the size of an earthquake is the amount of energy released by theearthquake. The amount of energy released is related to the Richter Scale by the following

equation:

Log E = 11.8 + 1.5 M

Where Log refers to the logarithm to the base 10, E is the energy released in ergs, and M is the

Richter Magnitude.

Anyone with a hand calculator can solve this equation by plugging in various values of M and

solving for E, the energy released. I've done the calculation for you in the following table:

Richter Magnitude Energy

(ergs)

Factor

1 2.0 x 1031 x

2 6.3 x 10

3 2.0 x 1031 x

4 6.3 x 10

5 2.0 x 1031 x

6 6.3 x 10

7 2.0 x 1031 x

8 6.3 x 10


24/32

From these calculations you can see that each increase in 1 in Richter Magnitude

represents a 31 fold increase in the amount of energy released. Thus, a magnitude 7

earthquake releases 31 times more energy than a magnitude 6 earthquake. A magnitude8 earthquake releases 31 x 31 or 961 times more energy than a magnitude 6 earthquake.

The Hiroshima atomic bomb released an amount of energy equivalent to a magnitude5.5 earthquake. The largest earthquake recorded, the Alaska earthquake in 1964, had a

Richter Magnitude of about 8.6. Note that larger earthquakes are possible, but have not

been recorded by humans.

Earthquake Risk

The risk that an earthquake will occur close to where you live depends on whether or not

tectonic activity that causes deformation is occurring within the crust of that area. For the U.S.,

the risk is greatest in the most tectonically active area, that is near the plate margin in theWestern U.S. Here, the San Andreas Fault which forms the margin between the Pacific Plate

and the North American Plate, is responsible for about 1 magnitude 8 or greater earthquake percentury. Also in the western U.S. is the Basin and Range Province, where extensional stressesin the crust have created many normal faults that are still active. Historically, large earthquakes

have also occurred in the area of New Madrid, Missouri; Charleston, South Carolina; and an

area extending from New Jersey to Massachusetts. (See figure 10.10 in your text). Whyearthquakes occur in these other areas is not well understood. If earthquakes have occurred

before, they are expected to occur again.

Earthquake Damage

Many seismologists have said that "earthquakes don't kill people, buildings do". This is because

most deaths from earthquakes are caused by buildings or other human construction falling

down during an earthquake. Earthquakes located in isolated areas far from human populationrarely cause any deaths. Thus, in earthquake prone areas like California, there are strict

building codes requiring the design and construction of buildings and other structures that willwithstand a large earthquake. While this program is not always completely successful, one fact

stands out to prove its effectiveness. In 1986 an earthquake near San Francisco, California with

a Richter Magnitude of 7.1 killed about 40 people. Most were killed when a double deckedfreeway collapsed. About 10 months later, an earthquake with magnitude 6.9 occurred in the

Armenia, where no earthquake proof building codes existed. The death toll in the latter

earthquake was about 25,000!

Damage from earthquakes can be classified as follows:

Ground Shaking - Shaking of the ground caused by the passage of seismic waves nearthe epicenter of the earthquake is responsible for the collapse of most structures. The

intensity of ground shaking depends on distance from the epicenter and on the type of

bedrock underlying the area.

o In general, loose unconsolidated sediment is subject to more intense shaking


25/32

than solid bedrock.

o Damage to structures from shaking depends on the type of construction.

Concrete and masonry structures, because they are brittle are more susceptible

to damage than wood and steel structures, which are more flexible.

Ground Rupture - Ground rupture only occurs along the fault zone that moves duringthe earthquake. Thus structures that are built across fault zones may collapse, whereas

structures built adjacent to, but not crossing the fault may survive.

Fire - Fire is a secondary effect of earthquakes. Because power lines may be knocked

down and because natural gas lines may rupture due to an earthquake, fires are often

started closely following an earthquake. The problem is compounded if water lines arealso broken during the earthquake since there will not be a supply of water to extinguish

the fires once they have started. In the 1906 earthquake in San Francisco more than

90% of the damage to buildings was caused by fire.

Rapid Mass-Wasting Processes - In mountainous regions subjected to earthquakes

ground shaking may trigger rapid mass-wasting events like rock and debris falls, rockand debris slides, slumps, and debris avalanches.

Liquefaction -

Liquefactionis a

processes that occurs inwater-saturated

unconsolidated

sediment due to

shaking. In areasunderlain by such

material, the

groundshaking causesthe grains to loose grain

to grain contact, and

thus the material tends

to flow.

You can demonstrate this process to yourself next time your go the beach. Stand on the

sand just after an incoming wave has passed. The sand will easily support your weight

and you will not sink very deeply into the sand if you stand still. But, if you start toshake your body while standing on this wet sand, you will notice that the sand begins to

flow as a result of liquefaction, and your feet will sink deeper into the sand.

Tsunamis - Tsunamis are giant ocean waves that can rapidly travel across oceans, as we

discussed in the Oceans and Their Margins. Earthquakes that occur along coastal areascan generate tsunamis, which can cause damage thousands of kilometers away on the


26/32

other side of the ocean.

World Distribution of Earthquakes

The distribution of earthquakes is referred to as seismicity. Most earthquakes occur alongrelatively narrow belts that coincide with plate boundaries (see figure 10.15 in your text).

This makes sense, sinceplate boundaries are

zones along which

lithospheric plates mover

relative to one another.Earthquakes along these

zones can be divided into

shallow focus

earthquakes that havefocal depths less than

about 70 km and deep

focus earthquakes thathave focal depths

between 75 and 700 km.

Earthquakes at Diverging Plate Boundaries. Diverging plate boundaries are zones where

two plates move away from each other, such as at oceanic ridges. In such areas the

lithosphere is in a state of tensional stress and thus normal faults and rift valleys occur.

Earthquakes that occur along such boundaries show normal fault motion and tend to beshallow focus earthquakes, with focal depths less than about 20 km. Such shallow focal

depths indicate that the brittle lithosphere must be relatively thin along these divergingplate boundaries.

Earthquakes at Transform Fault Boundaries. Transform fault boundaries are plate

boundaries where lithospheric plates slide past one another in a horizontal fashion. TheSan Andreas Fault of California is one of the longer transform fault boundaries known.

Earthquakes along these boundaries show strike-slip motion on the faults and tend to be

shallow focus earthquakes with depths usually less than about 50 km.

Earthquakes at Converging Plate Boundaries - Convergent plate boundaries are

boundaries where two plates run into each other. Thus, they tend to be zones where

compressional stresses are active and thus reverse faults or thrust faults are common.There are two types of converging plate boundaries. (1) subduction boundaries, where

oceanic lithosphere is pushed beneath either oceanic or continental lithosphere; and (2)

collision boundaries where two plates with continental lithosphere collide.

o Subduction boundaries -At subduction boundaries cold oceanic lithosphere ispushed back down into the mantle where two plates converge at an oceanic

trench. Because the subducted lithosphere is cold it remains brittle as it descends


27/32

and thus can fracture under the compressional stress. When it fractures, it

generates earthquakes that define a zone of earthquakes with increasing focal

depths beneath the overriding plate. This zone of earthquakes is called theBenioff Zone. Focal depths of earthquakes in the Benioff Zone can reach down

to 700 km.

o Collision boundaries - At collisional boundaries two plates of continental

lithosphere collide resulting in fold-thrust mountain belts. Earthquakes occur

due to the thrust faulting and range in depth from shallow to about 200 km.

The Earth's Internal Structure

Much of what we know about the interior of the Earth comes from knowledge of seismic wavevelocities and their variation with depth in the Earth. Recall that body wave velocities are as

follows:

Vp= [(K + 4/3)/]Vs= [( )/]Where K = incompressibility

= rigidity

= density

If the properties of the earth, i.e. K, , and where the same throughout, then Vpand Vswouldbe constant throughout the Earth and seismic waves would travel along straight line pathsthrough the Earth. We know however that density must change with depth in the Earth, because


28/32

the density of the Earth is 5,200 kg/cubic meter and density of crustal rocks is about 2,500

kg/cubic meter. If the density were the only property to change, then we could make estimates

of the density, and predict the arrival times or velocities of seismic waves at any point awayfrom an earthquake. Observations do not follow the predictions, so, something else must be

happening. In fact we know that K, , and change due to changing temperatures, pressures

and compositions of material. The job of seismology is, therefore, to use the observed seismicwave velocities to determine how K, , and change with depth in the Earth, and then inferhow P, T, and composition change with depth in the Earth. In other words to tell us something

about the internal structure of the Earth.

Reflection and Refraction of Seismic Waves.

If composition (or physical properties) change abruptly at some interface, then seismic wave

will both reflect off the interface and refract (or bend) as they pass through the interface. Two

cases of wave refraction can be recognized.

1. If the seismic wavevelocity in the rock

above an interface is lessthan the seismic wave

velocity in the rock

below the interface, the

waves will be refracted

or bent upward relativeto their original path.

If the seismic wave velocity decreases when passing into the rock below the interface,

the waves will be refracted down relative to their original path.


29/32

If the seismic wave velocities

gradually increase with depth in

the Earth, the waves willcontinually be refracted along

curved paths that curve back

toward the Earth's surface.

One of the earliest

discoveries of

seismology was adiscontinuity at a depth

of 2900 km where the

velocity of P-wavessuddenly decreases. This

boundary is the boundary

between the mantle andthe core and was

discovered because of a

zone on the opposite side

of the Earth from an

Earthquake focusreceives no direct P-

waves because the P-

waves are refractedinward as a result of the

sudden decrease in

velocity at the boundary.This zone is called a P-


30/32

wave shadow zone.

This discovery was

followed by the

discovery of an S-waveshadow zone. The S-

wave shadow zoneoccurs because no S-

waves reach the area on

the opposite side of theEarth from the focus.

Since no direct S-wavesarrive in this zone, it

implies that no S-waves

pass through the core.

This further implies thevelocity of S-wave in the

core is 0. In liquids =

0, so S-wave velocity is

also equal to 0. From this

it is deduced that thecore, or at least part of

the core is in the liquidstate, since no S-waves

are transmitted through

liquids. Thus, the S-waveshadow zone is best

explained by a liquid

outer core.

Seismic Wave Velocities in the Earth

Over the years seismologists have collected data on how seismic wave velocities vary with

depth in the Earth. Distinct boundaries, called discontinuities are observed when there is suddenchange in physical properties or chemical composition of the Earth. From these discontinuities,

we can deduce something about the nature of the various layers in the Earth. As we discussed

way back at the beginning of the course, we can look at the Earth in terms of layers of differingchemical composition, and layers of differing physical properties.


31/32


32/32

the outer core must be liquid since S-wave velocities are 0. At a depth of about4800 km the sudden increase in P-wave velocities indicate a solid inner core.The core appears to have a composition consistent with mostly Iron with small

amounts of Nickel.

Layers of Different Physical Properties

o At a depth of about 100 km there is a sudden decrease in both P and S-wavevelocities. This boundary marks the base of the lithosphere and the top of the

asthenosphere. The lithosphere is composed of both crust and part of the upper

mantle. It is a brittle layer that makes up the plates in plate tectonics, and appears

to float and move around on top of the more ductile asthenosphere.

o At the top of the asthenosphere is a zone where both P- and S-wave velocitiesare low. This zone is called the Low-Velocity Zone (LVZ). It is thought that the

low velocities of seismic waves in this zone are caused by temperatures

approaching the partial melting temperature of the mantle, causing the mantle inthis zone to behave in a very ductile manner.

o At adepth of 400 km there is an abrupt increase in the velocities of seismicwaves, thus this boundary is known as the400 - Km D iscontinu ity. Experimentson mantle rocks indicate that this represents a temperature and pressure where

there is a polymorphic phase transition, involving a change in the crystal

structure of Olivine, one of the most abundant minerals in the mantle.

o Another abrupt increase in seismic wave velocities occurs at a depth of 670 km.

It is uncertain whether this discontinuity, known as the670 Km Discontinu ity, isthe result of a polymorphic phase transition involving other mantle minerals or a

compositional change in the mantle, or both.

Return to EENS 111 Page
http://www.tulane.edu/~sanelson/geol111/index.htmlhttp://www.tulane.edu/~sanelson/geol111/index.htmlhttp://www.tulane.edu/~sanelson/geol111/index.html

seismic and the earth

Documents