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Next Generation
Sunshine State Standards
Chapter 8
LA.910.2.2.3. The student will organize information to show understanding or relationships
among facts, ideas, and events (e.g., representing key points within text through charting,
mapping, paraphrasing, summarizing, comparing, contrasting, or outlining).
LA.910.4.2.2. The student will record information and ideas from primary and/or secondary
sources accurately and coherently, noting the validity and reliability of these sources and
attributing sources of information.
MA.912.S.3.2. Collect, organize, and analyze data sets, determine the best format for the
data and present visual summaries from the following:
bar graphs
line graphs
cumulative frequency (ogive) graphs
SC.912.E.6.1. Describe and differentiate the layers of Earth and the interactions among them.
SC.912.N.1.1. Define a problem based on a specific body of knowledge, for example:
biology, chemistry, physics, and earth/space science, and do the following:
1.
3.
4.
7.
8.
9.
pose questions about the natural world,
examine books and other sources of information to see what is already known,
review what is known in light of empirical evidence,
pose answers, explanations, or descriptions of events,
generate explanations that explicate or describe natural phenomena (inferences),
use appropriate evidence and reasoning to justify these explanations to others
SC.912.N.1.4. Identify sources of information and assess their reliability according to the strict
standards of scientific investigation.
SC.912.N.2.2. Identify which questions can be answered through science and which questions
are outside the boundaries of scientific investigation, such as questions addressed by other
ways of knowing, such as art, philosophy, and religion.
Florida Sunshine State Standards Chapter 8
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Overall Instructional Quality
The major tool introduces and builds science concepts as a coherent whole. It provides
opportunities to students to explore why a scientific idea is important and in which contexts
that a science idea can be useful. In other words, the major tool helps students learn
the science concepts in depth. Additionally, students are given opportunities to connect
conceptual knowledge with procedural knowledge and factual knowledge. Overall, there
is an appropriate balance of skill development and conceptual understanding.
Tasks are engaging and interesting enough that students want to pursue them. Real world
problems are realistic and relevant to students’ lives.
Problem solving is encouraged by the tasks presented to students. Tasks require students
to make decisions, determine strategies, and justify solutions.
Students are given opportunities to create and use representations to organize, record,
and communicate their thinking.
Tasks promote use of multiple representations and translations among them. Students use
a variety of tools to understand a single concept.
The science connects to other disciplines such as reading, art, mathematics, and history.
Tasks represent scientific ideas as interconnected and building upon each other.
Benchmarks from the Nature of Science standard are both represented explicitly and
integrated throughout the materials.
Florida Sunshine State Standards Chapter 8
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Earthquakes and Earth’s Interior
C H A P T E R
8
Aftermath of a devastating tsunami at the coastal
town of Banda Aceh on
the Indonesian island of
Sumatra, December 26,
2004. (Photo by Mark
Pearson/Alamy)
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O
n October 17, 1989, at 5:04 P.M. Pacific daylight time, millions of television viewers
around the world were settling in to watch the third game of the World Series. Instead, they saw their television sets go black as tremors hit San Francisco’s Candlestick Park. Although the earthquake was centered in a remote section of the Santa Cruz
Mountains, 100 kilometers to the south, major damage occurred in the Marina District of
San Francisco.
The most tragic result of the violent shaking was the collapse of some double-decked
sections of Interstate 880, also known as the Nimitz Freeway. The ground motions caused
the upper deck to sway, shattering the concrete support columns along a mile-long section
of the freeway. The upper deck then collapsed onto the lower roadway, flattening cars as if
they were aluminum beverage cans. This earthquake, named the Loma Prieta quake for its
point of origin, claimed 67 lives.
In mid-January 1994, less than five years after the Loma Prieta earthquake devastated
portions of the San Francisco Bay Area, a major earthquake struck the Northridge area of
Los Angeles. Although not the fabled “Big One,” this moderate 6.7-magnitude earthquake
left 57 dead, more than 5,000 injured, and tens of thousands of households without water
and electricity. The damage exceeded $40 billion and was attributed to an apparently unknown fault that ruptured 18 kilometers (11 miles) beneath Northridge.
The Northridge earthquake began at 4:31 A.M. and lasted roughly 40 seconds. During
this brief period, the quake terrorized the entire Los Angeles area. In the three-story Northridge Meadows apartment complex, 16 people died when sections of the upper floors collapsed onto the first-floor units. Nearly 300 schools were seriously damaged, and a dozen
major roadways buckled. Among these were two of California’s major arteries—the Golden
State Freeway (Interstate 5), where an overpass collapsed completely and blocked the
roadway, and the Santa Monica Freeway. Fortunately, these roadways had practically no
traffic at this early morning hour.
In nearby Granada Hills, broken gas lines were set ablaze while the streets flooded from
broken water mains. Seventy homes burned in the Sylmar area. A 64-car freight train derailed, including some cars carrying hazardous cargo. But it is remarkable that the destruction was not greater. Unquestionably, the upgrading of structures to meet the requirements
of building codes developed for this earthquake-prone area helped minimize what could
have been a much greater human tragedy.
Students Sometimes Ask . . .
How often do earthquakes occur?
All the time—in fact, there are literally thousands of earthquakes
daily! Fortunately, the majority of them are too small to be felt by
people, and many of them occur in remote regions. Their existence is known only because of sensitive seismographs.
220
What Is an Earthquake?
Forces Within
Earthquakes
An earthquake is the vibration of Earth produced by the
rapid release of energy (Figure 8.1). Most often earthquakes
are caused by slippage along a fault in Earth’s crust. The energy released radiates in all directions from its source, the
focus 1foci = a point2, in the form of waves. These waves
are analogous to those produced when a stone is dropped
into a calm pond (Figure 8.2). Just as the impact of the stone
sets water waves in motion, an earthquake generates seismic
waves that radiate throughout the Earth. Even though the
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FIGURE 8.1 Destruction caused by a major earthquake that struck northwestern Turkey on August
17, 1999. More than 17,000 people perished. (Photo by Yann Arthus-Bertrand/Peter Arnold, Inc.)
energy dissipates rapidly with increasing distance from the
focus, sensitive instruments located throughout the world
record the event.
More than 30,000 earthquakes that are strong enough to
be felt occur worldwide annually. Fortunately, most are minor
tremors and do very little damage. Only about 75 significant
earthquakes take place each year, and many of these occur
in remote regions. Occasionally, however, a large earthquake
occurs near a large population center. Under these conditions,
an earthquake is among the most destructive natural forces
on Earth.
The shaking of the ground, coupled with the liquefaction
of some soils, wreaks havoc on buildings and other structures. In addition, when a quake occurs in a populated area,
power and gas lines are often ruptured, causing numerous
fires. In the famous 1906 San Francisco earthquake, much of
the damage was caused by fires (Figure 8.3), which quickly
became uncontrollable when broken water mains left firefighters with only trickles of water.
Earthquakes and Faults
Epicenter
Wave
fronts
Focus
Fault
FIGURE 8.2 The focus of an earthquake is located at depth. The
surface location directly above it is called the epicenter.
The tremendous energy released by atomic explosions or by
volcanic eruptions can produce an earthquake, but these
events are relatively weak and infrequent. What mechanism
produces a destructive earthquake? Ample evidence exists
that Earth is not a static planet. We know that Earth’s crust
has been uplifted at times, because we have found numerous
ancient wavecut benches many meters above the level of the
highest tides. Other regions exhibit evidence of extensive subsidence. In addition to these vertical displacements, offsets in
fence lines, roads, and other structures indicate that horizontal movement is common (Figure 8.4). These movements are
associated with large fractures in Earth’s crust called faults.
Typically, earthquakes occur along preexisting faults that
formed in the distant past along zones of weakness in Earth’s
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FIGURE 8.3 San Francisco in flames after the 1906 earthquake. (Reproduced from the collection
of the Library of Congress) Inset photo shows fire triggered when a gas line ruptured during the
Northridge earthquake in southern California in 1994. (AFP/Getty Images)
crust. Some are very large and can generate major earthquakes. One example is the San Andreas Fault, which is a
transform fault boundary that separates two great sections
of Earth’s lithosphere: the North American plate and the Pacific plate. Other faults are small and produce only minor
earthquakes.
Most faults are not perfectly straight or continuous; instead, they consist of numerous branches and smaller fractures that display kinks and offsets. Such a pattern is
displayed in Figure 10.A (p. 292), which shows that the San
Andreas Fault is actually a system that consists of several
large faults and innumerable small fractures (not shown).
It is also clear that most faults are locked, except for brief, abrupt movements that accomFIGURE 8.4 Slippage along a fault produced an offset in this orange grove east of
Calexico, California. (Photo by John S. Shelton) Inset photo shows a fence offset 2.5
pany an earthquake rupture. The primary
meters (8.5 feet) during the 1906 San Francisco earthquake. (Photo by G. K. Gilbert, U.S.
reason faults are locked is that the confining
Geological Survey)
pressure exerted by the overlying crust is enormous and essentially squeezes the fractures in
the crust shut. Nevertheless, even faults that
have been inactive for thousands of years can
rupture again if the stresses acting on the region
increase sufficiently.
Discovering the Cause
of Earthquakes
222
The actual mechanism of earthquake generation
eluded geologists until H. F. Reid of Johns Hopkins University conducted a study following the
great 1906 San Francisco earthquake. The earthquake was accompanied by horizontal surface
displacements of several meters along the northern portion of the San Andreas Fault. Field investigations determined that during this single
earthquake, the Pacific plate lurched as much as
4.7 meters (15 feet) northward relative to the adjacent North American plate.
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What Is an Earthquake?
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The mechanism for earthquake formation that Reid deduced
from this information is illustrated in Figure 8.5. Part A of the
figure shows an existing fault, or break in the rock. In part B, tectonic forces ever so slowly deform the crustal rocks on both
sides of the fault, as demonstrated by the bent features. Under
these conditions, rocks are bending and storing elastic energy,
much like a wooden stick does if bent. Eventually, the frictional
resistance holding the rocks in place is overcome. As slippage
occurs at the weakest point (the focus), displacement will exert
stress farther along the fault, where additional slippage will
occur, releasing the built-up strain (Figure 8.5C). This slippage
allows the deformed rock to “snap back.” The vibrations we
know as an earthquake occur as the rock elastically returns to
its original shape. The “springing back” of the rock was termed
elastic rebound by Reid, because the rock behaves elastically,
much like a stretched rubber band does when it is released.
Most of the motion along faults can be satisfactorily explained by the plate tectonics theory, which states that large
slabs of Earth’s lithosphere are in continual slow motion.
These mobile plates interact with neighboring plates, straining and deforming the rocks at their margins. In fact, it is
along faults associated with plate boundaries that most earthquakes occur. Furthermore, earthquakes are repetitive: As
soon as one is over, the continuous motion of the plates adds
strain to the rocks until they eventually fail again.
In summary, most earthquakes are produced by the rapid
release of elastic energy stored in rock that has been subjected
to great stress. Once the strength of the rock is exceeded, it
suddenly ruptures, causing the vibrations of an earthquake. EarthFIGURE 8.5 Elastic rebound. As rock is deformed it bends, storing elastic energy. Once the rock is
strained beyond its breaking point it ruptures, releasing the stored-up energy in the form of
quakes most often occur along
earthquake waves.
existing faults whenever the frictional forces on the fault surfaces
Deformation of rocks
Deformation of a limber stick
are overcome (see Box 8.1).
Foreshocks
and Aftershocks
Stream
Fault
A. Original position
A. Original position
Fault
B. Buildup of strain
B. Buildup of strain
C. Slippage (earthquake)
C. Slippage (earthquake)
Stream
offset
D. Strain released
D. Strain released
The intense vibrations of the 1906
San Francisco earthquake lasted
about 40 seconds. Although most
of the displacement along the
fault occurred in this rather short
period, additional movements
along this and other nearby faults
continued for several days following the main quake. The adjustments that follow a major
earthquake often generate smaller
earthquakes called aftershocks.
Although these aftershocks are
usually much weaker than the
main earthquake, they can sometimes destroy already badly
weakened structures. This occurred, for example, during a
1988 earthquake in Armenia. A
large aftershock of magnitude 5.8
collapsed many structures that
had been weakened by the main
tremor.
In addition, small earthquakes
called foreshocks often precede
a major earthquake by days or, in
some cases, by as much as several
years. Monitoring of these foreshocks has been used as a means
of predicting forthcoming major
earthquakes, with mixed success.
We consider the topic of earthquake prediction in a later section
of this chapter.
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BOX 8.1 PEOPLE AND THE ENVIRONMENT
Damaging Earthquakes East of the Rockies
When you think earthquake, you probably
think of California and Japan. However, six
major earthquakes have occurred in the central and eastern United States since colonial
times. Three of these had estimated Richter
magnitudes of 7.5, 7.3, and 7.8, and they
were centered near the Mississippi River
Valley in southeastern Missouri. Occurring
on December 16, 1811, January 23, 1812, and
February 7, 1812, these earthquakes, plus
numerous smaller tremors, destroyed the
town of New Madrid, Missouri, triggered
massive landslides, and caused damage
over a six-state area. The course of the Mississippi River was altered, and Tennessee’s
Reelfoot Lake was enlarged. The distances
over which these earthquakes were felt are
truly remarkable. Chimneys were reported
downed in Cincinnati, Ohio, and Richmond, Virginia, while Boston residents, located 1,770 kilometers (1,100 miles) to the
northeast, felt the tremor.
Despite the history of the New Madrid
earthquake, Memphis, Tennessee, the
largest population center in the area today,
does not have adequate earthquake provisions in its building code. Furthermore,
because Memphis is located on unconsolidated floodplain deposits, buildings are
more susceptible to damage than similar
structures built on bedrock. It has been estimated that if an earthquake the size of the
1811–1812 New Madrid event were to strike
in the next decade, it would result in casualties in the thousands and damages in tens
of billions of dollars.
Damaging earthquakes that occurred in
Aurora, Illinois (1909), and Valentine, Texas
(1931), remind us that other areas in the central United States are vulnerable.
The greatest historical earthquake in the
eastern states occurred August 31, 1886, in
Charleston, South Carolina. The event,
which spanned 1 minute, caused 60 deaths,
numerous injuries, and great economic loss
FIGURE 8.A Damage to Charleston, South Carolina, caused by the August 31, 1886,
earthquake. Damage ranged from toppled chimneys and broken plaster to total collapse.
(Photo courtesy of U.S. Geological Survey)
within a radius of 200 kilometers (120 miles)
of Charleston. Within 8 minutes, effects
were felt as far away as Chicago and St.
Louis, where strong vibrations shook the
upper floors of buildings, causing people to
rush outdoors. In Charleston alone, over 100
buildings were destroyed, and 90 percent of
the remaining structures were damaged
(Figure 8.A).
Numerous other strong earthquakes
have been recorded in the eastern United
States. New England and adjacent areas
have experienced sizable shocks since colonial times. The first reported earthquake in
the Northeast took place in Plymouth,
Massachusetts, in 1683, and was followed
in 1755 by the destructive Cambridge, Massachusetts, earthquake. Moreover, ever since
records have been kept, New York State
alone has experienced over 300 earthquakes
large enough to be felt.
San Andreas Fault:
An Active Earthquake Zone
The San Andreas is undoubtedly the most studied fault system in the world (Figure 8.6). Over the years, investigations
have shown that displacement occurs along discrete segments
that are 100 to 200 kilometers long. Furthermore, each fault
segment behaves somewhat differently from the others. Some
portions of the San Andreas exhibit a slow, gradual dis-
Earthquakes in the central and eastern
United States occur far less frequently than
in California. Yet history indicates that the
East is vulnerable. Furthermore, these
shocks east of the Rockies have generally
produced structural damage over a larger
area than counterparts of similar magnitude
in California. The reason is that the underlying bedrock in the central and eastern
United States is older and more rigid. As a
result, seismic waves are able to travel
greater distances with less weakening than
in the western United States. It is estimated
that for earthquakes of similar magnitude,
the region of maximum ground motion in
the East may be up to 10 times larger than
in the West. Consequently, the higher rate
of earthquake occurrence in the western
United States is balanced somewhat by the
fact that central and eastern U.S. quakes can
damage larger areas.
placement known as fault creep, which occurs relatively
smoothly and therefore with little noticable seismic activity.
Other segments regularly slip, producing small earthquakes.
Still other segments remain locked and store elastic energy
for hundreds of years before rupturing in great earthquakes.
The latter process is described as stick-slip motion, because
the fault exhibits alternating periods of locked behavior followed by sudden slippage. It is estimated that great earthquakes should occur about every 50 to 200 years along those
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Seismology: The Study of Earthquake Waves
225
still active. Currently, laser beams are used to measure the
relative motion between the opposite sides of this fault. These
measurements reveal a displacement of 2–5 centimeters (1–2
inches) per year. Although this seems slow, it produces substantial movement over millions of years.
To illustrate, in 30 million years this rate of displacement
would slide the western portion of California northward so
that Los Angeles, on the Pacific plate, would be adjacent to
San Francisco on the North American plate! More important
in the short term, a displacement of just 2 centimeters each
year produces 2 meters of offset every 100 years. Consequently, the 4 meters of displacement produced during the
1906 San Francisco earthquake should occur at least every
200 years along this segment of the fault zone. This fact lies
behind California’s concern for making buildings earthquakeresistant in anticipation of the inevitable “Big One.”
Seismology: The Study
of Earthquake Waves
Forces Within
Earthquakes
FIGURE 8.6 Trace of the San Andreas Fault, north of Landers California.
(Photo by Roger Ressmeyer/CORBIS)
sections of the San Andreas Fault that exhibit stick-slip motion. This knowledge is useful when assigning a potential
earthquake risk to a given segment of the fault zone.
The tectonic forces along the San Andreas Fault zone that
were responsible for the 1906 San Francisco earthquake are
Students Sometimes Ask . . .
Do moderate earthquakes decrease the chances of a major
quake in the same region?
No. This is due to the vast increase in release of energy associated
with higher-magnitude earthquakes. For instance, an earthquake
with a magnitude of 8.5 releases millions of times more energy
than the smallest earthquakes felt by humans. Similarly, thousands of moderate tremors would be needed to release the huge
amount of energy equal to one “great” earthquake.
The study of earthquake waves, seismology 1seismos =
shake, ology = the study of2, dates back to attempts by the
Chinese almost 2,000 years ago to determine the direction of
the source of each earthquake. Modern seismographs
1seismos = shake, graph = write2 are instruments that record
earthquake waves. Their principle is simple: A weight is freely
suspended from a support that is attached to bedrock (Figure
8.7). When waves from an earthquake reach the instrument,
the inertia of the weight keeps it stationary, while Earth and
the support vibrate. The movement of Earth in relation to the
stationary weight is recorded on a rotating drum. (Inertia is
the tendency of a stationary object to remain still, or a moving object to stay in motion.)
Modern seismographs amplify and record ground motion,
producing a trace as shown in Figure 8.8. These records,
called seismograms 1seismos = shake, gramma = what is
written2, reveal that seismic waves are elastic energy. This
energy radiates outward in all directions from the focus, as
you saw in Figure 8.2. The transmission of this energy can be
compared to the shaking of gelatin in a bowl that is jarred.
Seismograms reveal that two main types of seismic waves
are generated by the slippage of a rock mass. Some travel
along Earth’s outer layer and are called surface waves. Others travel through Earth’s interior and are called body waves.
Body waves are further divided into primary waves
(P waves) and secondary waves (S waves).
Body waves are divided into P and S waves by their mode
of travel through intervening materials. P waves are
push–pull waves—they push (compress) and pull (expand)
rocks in the direction the wave is traveling (Figure 8.9A).
Imagine holding someone by the shoulders and shaking
them. This push–pull movement is how P waves move
through the Earth. This wave motion is analogous to that generated by human vocal cords as they move air to create
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Earthquakes and Earth’s Interior
Figure 8.9C. Unlike P waves, which
temporarily change the volume of the
intervening material by alternately
compressing and expanding it, S
waves temporarily change the shape
Bedrock
of the material that transmits them
(Figure 8.9D). Because fluids (gases
Weight hinged
and liquids) do not respond elastically
to allow
movement
to changes in shape, they will not transmit S waves.
The motion of surface waves is
somewhat more complex. As surface
waves travel along the ground, they
Support
Mass does not
moves with
cause the ground and anything resting
move with
Earth
B.
ground motion
upon it to move, much like ocean
due to inertia
Pen
swells toss a ship. In addition to their
up-and-down motion, surface waves
have a side-to-side motion similar to
an S wave oriented in a horizontal
Rotating drum
records
motion
plane. This latter motion is particularly
Earth moves
damaging to the foundations of
structures.
Bedrock
By observing a typical seismic
record, as shown in Figure 8.8, you can
A.
see a major difference among these
FIGURE 8.7 Principle of the seismograph. The inertia of the suspended mass tends to keep it
seismic waves: P waves arrive at the
motionless, while the recording drum, which is anchored to bedrock, vibrates in response to seismic
recording station first, then S waves,
waves. Thus, the stationary mass provides a reference point from which to measure the amount of
displacement occurring as the seismic waves pass through the ground below.
and then surface waves. This is a con(Photo by Zephyr/Photo Rearchers, Inc.)
sequence of their speeds. To illustrate,
the velocity of P waves through gransound. Solids, liquids, and gases resist a change in volume
ite within the crust is about 6 kilometers (4 miles) per second.
when compressed and will elastically spring back once the
Under the same conditions, S waves travel at 3.5 kilometers
force is removed (Figure 8.9B). Therefore, P waves, which are
(2 miles) per second. Differences in density and elastic propcompressional waves, can travel through all these materials.
erties of the rock greatly influence the velocities of these
In contrast, S waves shake the particles at right angles to
waves. Generally, in any solid material, P waves travel about
their direction of travel. This can be illustrated by fastening
1.7 times faster than S waves, and surface waves can be exone end of a rope and shaking the other end, as shown in
pected to travel at 90 percent of the velocity of the S waves.
As you will see, seismic waves allow us to determine the location and magnitude of earthquakes. In addition, seismic waves
FIGURE 8.8 Typical seismic record. Note the time interval (about 5
provide us with an important tool for probing Earth’s interior.
minutes) between the arrival of the first P wave and the arrival of the first
S wave.
Locating an Earthquake
Surface waves
First P wave
First S wave
1 minute
(Earlier)
T I M E
(Later)
Forces Within
A Earthquakes
Recall that the focus is the place within Earth where earthquake waves originate. The epicenter 1epi = upon, center =
a point2 is the location on the surface directly above the focus
(see Figure 8.2).
The difference in velocities of P and S waves provides a
method for locating the epicenter. The principle used is analogous to a race between two autos, one faster than the other.
The P wave always wins the race, arriving ahead of the S
wave. But the greater the length of the race, the greater will
be the difference in the arrival times at the finish line (the seismic station). Therefore, the greater the interval measured on
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Locating an Earthquake
227
Slinky at rest
Push slinky
Compress
Wave direction
Expand Compress
Wave direction
Particle motion
A. P waves generated using a slinky
B. P waves traveling along the surface
Rope at rest
Shake rope
Particle motion
Wave direction
Wave direction
Particle motion
C. S waves generated using a rope
D. S waves traveling along the surface
FIGURE 8.9 Types of seismic waves and their characteristic motion. (Note that during a strong
earthquake, ground shaking consists of a combination of various kinds of seismic waves.) A. As
illustrated by a slinky, P waves are compressional waves that alternately compress and expand the
material through which they pass. B. The back-and-forth motion produced as compressional waves
travel along the surface can cause the ground to buckle and fracture, and may cause power lines to
break. C. S waves cause material to oscillate at right angles to the direction of wave motion. D.
Because S waves can travel in any plane, they produce up-and-down and sideways shaking of the
a seismogram between the arrival of the first P wave and the
first S wave, the greater the distance to the earthquake source.
A system for locating earthquake epicenters was developed
by using seismograms from earthquakes whose epicenters
could be easily pinpointed from physical evidence. From these
seismograms, travel-time graphs were constructed (Figure
8.10B). The first travel-time graphs were greatly improved when
seismograms became available from nuclear explosions, because the precise location and time of detonation were known.
Using the sample seismogram in Figure 8.10A and the
travel-time curve in Figure 8.10B, we can determine the distance separating the recording station from the earthquake
in two steps: (1) using the seismogram, determine the time interval between the arrival of the first P wave and the first S
wave, and (2) using the travel-time graph, find the P-S interval on the vertical axis and use that information to determine
the distance to the epicenter on the horizontal axis. From this
information, we can determine that this earthquake occurred
3800 kilometers (2350 miles) from the recording instrument.
Now we know the distance, but what about the direction?
The epicenter could be in any direction from the seismic sta-
tion. As shown in Figure 8.11, the precise location of a quake
can be found when the distance is known from three or more
different seismic stations. On a globe, a circle is drawn around
each station with each circle’s radius equal to the station’s
distance from the epicenter. The point where the three circles
intersect is the epicenter of the quake. This method is called
triangulation.
About 95 percent of the energy released by earthquakes
originates in a few relatively narrow zones (Figure 8.12). The
greatest energy is released along a path around the outer edge
of the Pacific Ocean known as the circum-Pacific belt. Included
in this zone are regions of great seismic activity, such as Japan,
the Philippines, Chile, and numerous volcanic island chains,
as exemplified by Alaska’s Aleutian Islands.
Figure 8.12 reveals another continuous belt that extends
for thousands of kilometers through the world’s oceans. This
zone coincides with the oceanic ridge system, an area of frequent but low-intensity seismic activity. By comparing this
figure with Figure 7.10 (pp. 196–197), you can see a close correlation between the location of earthquake epicenters and
plate boundaries.
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P-wave
First
S-wave
km
00
84
Time interval
5 minutes
0
1
2
3
4
5
Paris
Montréal
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Travel time in minutes
A. Seismogram
Epicenter
Distance in miles
500
1000
1500
2000
2500
3000
15
0k
550
m
São Paulo
14
13
Time in minutes
12
11
ve
wa
S
10
e
rv
cu
5 min.
time
interval
9
8
7
ve
wa
P-
6
5
rve
cu
4
Intensity Scales
3
2
1
0
FIGURE 8.11 An earthquake epicenter is located using the distances
obtained from three seismic stations. On a globe, a circle is drawn around
each station, with each circle’s radius equal to the station’s distance from
the earthquake’s epicenter. Today, computers using data from numerous
seismic stations can rapidly pinpoint large earthquakes.
1000
2000
3000
Distance in kilometers
4000
B. Travel-time graph
FIGURE 8.10 Using a seismogram and a travel-time graph to determine
the distance to an earthquake’s epicenter. A. The time delay between the
arrival of the first P- and S-waves on this seismogram is 5 minutes.
B. Using the travel-time graph, it is determined that the epicenter is
roughly 3,800 kilometers (2,350 miles) from the seismic station.
Measuring the Size of Earthquakes
Seismologists have employed a variety of methods to obtain
two fundamentally different measures that describe the size
of an earthquake: intensity and magnitude. The first of these
to be used was intensity—a measure of the degree of earthquake shaking at a given locale based on the amount of damage. With the development of seismographs, it became clear
that a quantitative measure of an earthquake based on seismic records rather than uncertain personal estimates of damage was desirable. The measurement that was developed,
called magnitude, relies on calculations that use data provided by seismic records (and other techniques) to estimate
the amount of energy released at the source of the earthquake.
Until a little more than a century ago, historical records provided the only accounts of the severity of earthquake shaking and destruction. Using these descriptions—which were
compiled without any established standards for reporting—
made accurate comparisons of earthquake sizes difficult, at
best.
In order to standardize the study of earthquake severity,
scientists developed various intensity scales that considered
damage done to buildings, as well as individual descriptions
of the event, and secondary effects such as landslides and the
extent of ground rupture. By 1902, Guiseppe Mercalli had developed a relatively reliable intensity scale, which in a modified form is still used today. The Modified Mercalli Intensity
Scale shown in Table 8.1 was developed using California
buildings as its standard, but it is appropriate for use throughout most of the world to estimate the strength of an earthquake (Figure 8.13). For example, if some well-built wood
structures and most masonry buildings are destroyed by an
earthquake, a region would be assigned an intensity of X on
the Mercalli scale (Table 8.1).
Despite their usefulness in providing seismologists with
a tool to compare earthquake severity, particularly in regions
where there are no seismographs, intensity scales have severe drawbacks. In particular, intensity scales are based on
effects (largely destruction) of earthquakes that depend not
only on the severity of ground shaking but also on factors
such as population density, building design, and the nature
of surface materials.
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FIGURE 8.12 Distribution of the 14,229 earthquakes with magnitudes equal to or greater than 5 for
a 10-year period. Inset photo shows earthquake damage near Ismit, Turkey, in 1999. (Photo courtesy
of CORBIS/SYGMA)
Magnitude Scales
In order to compare earthquakes across the globe, a measure
was needed that did not rely on parameters that vary considerably from one part of the world to another, such as construction practices. As a consequence, a number of magnitude
scales were developed.
Richter Magnitude In 1935, Charles Richter of the California Institute of Technology developed the first magnitude
scale using seismic records to estimate the relative sizes of
earthquakes. As shown in Figure 8.14 (top), the Richter scale
is based on the amplitude of the largest seismic wave (P, S, or
surface wave) recorded on a seismogram. Because seismic
waves weaken as the distance between the earthquake focus
and the seismograph increases (in a manner similar to light),
TABLE 8.1
I
II
III
IV
V
VI
VII
VIII
IX
X
XI
XII
Richter developed a method that accounted for the decrease
in wave amplitude with increased distance. Theoretically, as
long as the same, or equivalent, instruments were used, monitoring stations at various locations would obtain the same
Richter magnitude for every recorded earthquake. (Richter
selected the Wood-Anderson seismograph as the standard
recording device.)
Although the Richter scale has no upper limit, the largest
magnitude recorded on a Wood-Anderson seismograph was
8.9. These great shocks release approximately 1026 ergs of energy—roughly equivalent to the detonation of 1 billion tons
of TNT. Conversely, earthquakes with a Richter magnitude of
less than 2.0 are not felt by humans. With the development of
more sensitive instruments, tremors of a magnitude of minus
2 were recorded. Table 8.2 shows how Richter magnitudes
and their effects are related.
Modified Mercalli Intensity Scale
Not felt except by a very few under especially favorable circumstances.
Felt only by a few persons at rest, especially on upper floors of buildings.
Felt quite noticeably indoors, especially on upper floors of buildings, but many people do not recognize it as an earthquake.
During the day felt indoors by many, outdoors by few. Sensation like heavy truck striking building.
Felt by nearly everyone, many awakened. Disturbances of trees, poles, and other tall objects sometimes noticed.
Felt by all; many frightened and run outdoors. Some heavy furniture moved; few instances of fallen plaster or damaged chimneys. Damage
slight.
Everybody runs outdoors. Damage negligible in buildings of good design and construction; slight-to-moderate in well-built ordinary structures;
considerable in poorly built or badly designed structures.
Damage slight in specially designed structures; considerable in ordinary substantial buildings with partial collapse; great in poorly built
structures. (Fall of chimneys, factory stacks, columns, monuments, walls.)
Damage considerable in specially designed structures. Buildings shifted off foundations. Ground cracked conspicuously.
Some well-built wooden structures destroyed. Most masonry and frame structures destroyed. Ground badly cracked.
Few, if any, (masonry) structures remain standing. Bridges destroyed. Broad fissures in ground.
Damage total. Waves seen on ground surfaces. Objects thrown upward into air.
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a
N
of
Jap
an
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e
A
S
Epicenter
III IV
V VI
A
J
II
P
FIGURE 8.13 Zones of destruction associated with an earthquake that
struck Japan in 1925. Intensity levels based on the Modified Mercalli
Intensity Scale.
FIGURE 8.14 Illustration showing how the Richter magnitude of an
earthquake can be determined graphically using a seismograph record
from a Wood-Anderson instrument. First, measure the height (amplitude)
of the largest wave on the seismogram (23 mm) and then the distance to
the focus using the time interval between S and P waves (24 seconds).
Next, draw a line between the distance scale (left) and the wave
amplitude scale (right). By doing this, you should obtain the Richter
magnitude 1ML2 of 5. (Data from California Institute of Technology)
Amplitude
23 mm
24 sec.
P
S
0
10 20
Distance, S-P,
km
sec.
500
50
400
40
300
30
200
100
60
40
20
0.5
20
10
Seismograph record
Time
sec.
30 mm
Magnitude,
ML
Amplitude,
mm
100
6
50
5
20
4
5
10
8
6
3
2
4
2
0.5
1
1.2
10
20
1
2
0.1
0
Earthquakes vary enormously in strength, and great earthquakes produce wave amplitudes that are thousands of times
larger than those generated by weak tremors. To accommodate this wide variation, Richter used a logarithmic scale to express magnitude, where a tenfold increase in wave amplitude
corresponds to an increase of 1 on the magnitude scale. Thus,
the amount of ground shaking for a 5-magnitude earthquake
is 10 times greater than that produced by an earthquake having a Richter magnitude of 4.
In addition, each unit of Richter magnitude equates to
roughly a 32-fold energy increase. Thus, an earthquake with a
magnitude of 6.5 releases 32 times more energy than one with
a magnitude of 5.5, and roughly 1,000 times more energy than
a 4.5-magnitude quake. A major earthquake with a magnitude of 8.5 releases millions of times more energy than the
smallest earthquakes felt by humans.
Richter’s original goal was modest in that he only attempted to rank the earthquakes of southern California
(shallow-focus earthquakes) into groups of large, medium,
and small magnitude. Hence, Richter magnitude was designed to study nearby (or local) earthquakes and is denoted
by the symbol 1ML2 where M is for magnitude and L is for local.
The convenience of describing the size of an earthquake
by a single number that could be calculated quickly from seismograms makes the Richter scale a powerful tool. Furthermore, unlike intensity scales that could only be applied to
populated areas of the globe, Richter magnitudes could be
assigned to earthquakes in more remote regions and even to
events that occurred in the ocean basins. As a result, the
method devised by Richter was adapted to a number of different seismographs located throughout the world. In time,
seismologists modified Richter’s work and developed new
magnitude scales.
Moment Magnitude Seismologists have recently been employing a more precise measure called moment magnitude
1MW2, which can be calculated using several techniques. In
one method, the moment magnitude is calculated from field
studies using a combination of factors that include the average amount of displacement along the fault, the area of the
rupture surface, and the shear strength of the faulted rock—
a measure of how much energy a rock can store before it
suddenly slips and releases this energy in the form of an
earthquake (and heat).
The moment magnitude can also be readily calculated
from seismograms by examining very long period seismic
waves. The values obtained have been calibrated so that
small- and moderate-sized earthquakes have moment magnitudes that are roughly equivalent to Richter magnitudes.
However, moment magnitudes are much better for describing very large earthquakes. For example, on the moment
magnitude scale, the 1906 San Francisco earthquake, which
had a Richter magnitude of 8.3, would be demoted to 7.9 on
the moment magnitude scale, whereas the 1964 Alaskan
earthquake with an 8.3 Richter magnitude would be increased to 9.2. The strongest earthquake on record is the 1960
Chilean earthquake with a moment magnitude of 9.5.
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TABLE 8.2
231
Earthquake Magnitudes and Expected World Incidence
Richter Magnitudes
62.0
2.0–2.9
3.0–3.9
4.0–4.9
5.0–5.9
6.0–6.9
7.0–7.9
8.0 and above
Effects Near Epicenter
Generally not felt, but recorded
Potentially perceptible
Felt by some
Felt by most
Damaging shocks
Destructive in populous regions
Major earthquakes; inflict serious damage
Great earthquakes; destroy communities near epicenter
Estimated Number per Year
600,000
300,000
49,000
6,200
800
266
18
1.4
Source: Earthquake Information Bulletin and others.
Moment magnitude has gained wide acceptance among
seismologists and engineers because (1) it is the only magnitude scale that estimates adequately the size of very large
earthquakes; (2) it is a measure that can be derived mathematically from the size of the rupture surface and the amount
of displacement and it better reflects the total energy released
during an earthquake; and (3) it can be verified by two independent methods—field studies that are based on measurements of fault displacement and seismographic methods
using long-period waves.
Destruction from Earthquakes
Students Sometimes Ask . . .
I’ve heard that the safest place to be in a house during an
earthquake is in a doorframe. Is that really the best place?
No! An enduring earthquake image of California is a collapsed
adobe home with the door frame as the only standing part. From
this came the belief that a doorway is the safest place to be during an earthquake. In modern homes, doorways are no stronger
than any other part of the house and usually have doors that will
swing and can injure you.
If you’re inside, the best advice is to duck, cover, and hold. When
you feel an earthquake, duck under a desk or sturdy table. Stay
away from windows, bookcases, file cabinets, heavy mirrors,
hanging plants, and other heavy objects that could fall. Stay
under cover until the shaking stops. And, hold on to the desk or
table: If it moves, move with it.
The most violent earthquake to jar North America this century—the Good Friday Alaskan Earthquake—occurred in
1964. Felt throughout the state, the earthquake had a moment magnitude of 9.2 and reportedly lasted 3–4 minutes.
This event left 131 people dead, thousands homeless, and
the economy of the state badly disrupted because it occurred
near major towns and seaports (Figure 8.15). Had the
schools and business districts been open on this holiday, the
toll surely would have been
higher. Within 24 hours of the
FIGURE 8.15 Region most affected by the Good Friday earthquake of 1964. Note the location of
initial shock, 28 aftershocks
the epicenter (red dot). Inset photo shows the collapse of a street in Anchorage, Alaska, caused by
this earthquake. (Photo courtesy of U.S. Geological Survey)
were recorded, 10 of which
exceeded a Richter magnitude
of 6.
Damage from
Seismic Vibrations
The 1964 Alaskan earthquake
provided geologists with new
insights into the role of ground
shaking as a destructive force.
As the energy released by
an earthquake travels along
Earth’s surface, it causes the
ground to vibrate in a complex
manner by moving up and
down as well as from side to
side. The amount of structural
Anchorage
Whittier
Seward
Valdez
Chenega
Cook
Inlet
Gulf of Alaska
Kodiak
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Most large structures in Anchorage were damaged, even
though they were built according to the earthquake provisions of the Uniform Building Code. Perhaps some of that
destruction can be attributed to the unusually long duration
of this earthquake. Most quakes consist of tremors that last
less than a minute. The 1994 Northridge earthquake was felt
for about 40 seconds, and the strong vibrations of the 1989
Loma Prieta earthquake lasted less than 15 seconds. But the
Alaska quake reverberated for 3 to 4 minutes.
FIGURE 8.16 Damage to the five-story J.C. Penney Co. building,
Anchorage, Alaska. Very little structural damage was incurred by the
adjacent building. (Courtesy of NOAA)
damage attributable to the vibrations depends on several factors, including (1) the intensity and (2) duration of the vibrations, (3) the nature of the material upon which the structure
rests, and (4) the design of the structure.
All multistory structures in Anchorage were damaged by
the vibrations, but the more flexible wood-frame residential
buildings fared best. Figure 8.16 offers a striking example of
how construction variations affect earthquake damage; you
can see that the steel-frame building on the left withstood the
vibrations, whereas the poorly designed J.C. Penney building
was badly damaged. Engineers have learned that unreinforced masonry buildings are the most serious safety threat
in earthquakes.
Amplification of Seismic Waves Although the region within
20 to 50 kilometers (12 to 30 miles) of the epicenter will experience about the same intensity of ground shaking, the destruction varies considerably within this area. This difference
is mainly attributable to the nature of the ground on which
the structures are built. Soft sediments, for example, generally
amplify the vibrations more than solid bedrock. Thus, the
buildings located in Anchorage, which were situated on unconsolidated sediments, experienced heavy structural damage. By contrast, most of the town of Whittier, although much
nearer the epicenter, rests on a firm foundation of granite and
hence suffered much less damage. However, Whittier was
damaged by a seismic sea wave (described in the next
section).
Liquefaction In areas where unconsolidated materials are
saturated with water, earthquake vibrations can generate a
phenomenon known as liquefaction 1liqueo = to be fluid,
facio = to make2. Under these conditions, what had been a
stable soil turns into a mobile fluid that is not capable of supporting buildings or other structures (Figure 8.17). As a result, underground objects such as storage tanks and sewer
lines may literally float toward the surface of their newly liquefied environment. Buildings and other structures may settle and collapse. During the
1989 Loma Prieta earthquake, in
FIGURE 8.17 Effects of liquefaction. This tilted building rests on unconsolidated sediment that
San Francisco’s Marina District,
imitated quicksand during the 1985 Mexican earthquake. (Photo by James L. Beck)
foundations failed and geysers
of sand and water shot from the
ground, indicating that liquefaction had occurred (Figure
8.18).
What Is a Tsunami?
Large undersea earthquakes occasionally set in motion massive
waves of water called seismic
sea waves, or tsunami* 1tsu =
harbor, nami = waves2. These
destructive waves often are
called “tidal waves” by the
media. However, this name is
*“Seismic sea waves were given the name
tsunami by the Japanese, who have suffered a
great deal from them. The term tsunami is now
used worldwide.
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233
inappropriate, because these waves are not created by the
tidal effect of the Moon or Sun. Most tsunami result from vertical displacement along a fault located on the ocean floor, or
from a large underwater landslide triggered by an earthquake
(Figure 8.19).
Once formed, a tsunami resembles the ripples created
when a pebble is dropped into a pond. In contrast to ripples,
a tsunami advances across the ocean at speeds between 500
and 950 kilometers (300 and 600 miles) per hour. Despite this,
a tsunami in the open ocean can pass undetected because its
height is usually less than 1 meter and the distance between
wave crests is great, ranging from 100 to 700 kilometers (62
to 435 miles). Upon entering shallower coastal water, these
destructive waves are slowed and the water begins to pile up
to heights that occasionally exceed 30 meters (100 feet), as
shown in Figure 8.19. As the crest of a tsunami approaches
shore, it appears as a rapid rise in sea level with a turbulent
and chaotic surface.
The first warning of an approaching tsunami is usually a
rapid withdrawal of water from beaches. Some residents living near the Pacific basin have learned to heed this warning
and move to higher ground, because about 5 to 30 minutes
later the retreat of water is followed by a surge capable of extending hundreds of meters inland. In a successive fashion,
each surge is followed by a rapid oceanward retreat of the
water.
A.
B.
FIGURE 8.18 Liquefaction. A. These mud volcanoes were produced
by the Loma Prieta earthquake in 1989. Mud volcanoes form when
geysers of sand and water shoot from the ground, an indication that
liquefaction has occurred. (Photo by Richard Hilton, courtesy of Dennis
Fox) B. Students experiencing the “strength” of water-saturated soil, a
hallmark of liquefaction. (Photo by M. Miller)
Tsunami Damage from the 2004 Indonesian Earthquake A
massive undersea earthquake of moment magnitude 9.0 occurred near the island of Sumatra on December 26, 2004, and
sent waves of water racing across the Indian Ocean and Bay
of Bengal (Figure 8.20A). This tsunami was one of the deadliest natural disasters of any kind in modern times, claiming
more than 230,000 lives. As water surged several kilometers
inland, cars and trucks were flung around like toys in a bathtub, and fishing boats were rammed into homes (see chapter
opening photo). In some locations, the backwash of water
dragged bodies and huge amounts of debris out to sea.
FIGURE 8.19 Schematic drawing of a tsunami generated by displacement of the ocean floor. The
speed of a wave moving across the ocean correlates with ocean depth. As shown, waves moving in
deep water advance at speeds in excess of 800 kilometers per hour. Speed gradually slows to 50
kilometers per hour at depths of 20 meters. Decreasing depth slows the movement of the wave
column. As waves slow in shallow water, they grow in height until they topple and rush onto shore
with tremendous force. The size and spacing of these swells are not to scale.
Tsunami speed:
835 km/hr
Tsunami speed:
340 km/hr
Tsunami speed:
50 km/hr
Sea level
Water depth:
5500 meters
Subducti
ng pla
te
Displacement
Water depth:
900 meters
Water depth:
20 meters
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A.
B.
FIGURE 8.20 A massive earthquake 19.0 Mw2 off the Indonesian island of Sumatra sent a tsunami
racing across the Indian Ocean and Bay of Bengal on December 26, 2004. A. Unsuspecting foreign
tourists who walked seaward as the water receded, rush toward shore as the first of six tsunami rolled
toward Hat Rai Lay Beach near Krabi in southern Thailand. (AFP/Getty Images. Inc.) B. Tsunami
survivors walk among the debris from this earthquake-triggered event. (Photo by Kimmasa
Mayama/Reuters/CORBIS)
The destruction was indiscriminate, destroying luxury resorts and poor fishing hamlets on the Indian Ocean coast
(Figure 8.20B). Devastation was most severe along the southeast coast of Sri Lanka, in the Indonesian province of Aceh, in
the Indian state of Tamil Nadu, and on Thailand’s resort island of Phuket. Damages were reported as far away as the
Somalia coast of Africa, 4,100 kilometers (2,500 miles) west
of the earthquake epicenter.
The killer waves generated by this massive quake achieved
heights as great as 10 meters (33 feet) and struck many unprepared areas within three hours of the event. Although the
Pacific basin was equipped with deep-sea buoys and tide
gauges that can spot tsunami waves at sea, the Indian Ocean
was not. (The deep-sea buoys have pressure sensors that detect changes in pressure as the earthquake’s energy travels
through the ocean, and tide gauges measure the rise and fall
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Destruction from Earthquakes
in sea level.) The rarity of tsunami in the Indian Ocean also
contributed to the lack of preparedness for such an event. It
should come as no surprise that a tsunami warning system
for the Indian Ocean has now been established.
Tsunami Warning System In 1946, a large tsunami struck
the Hawaiian Islands without warning. A wave more than
15 meters (50 feet) high left several coastal villages in shambles. This destruction motivated the U.S. Coast and Geodetic
Survey to establish a tsunami warning system for coastal
areas of the Pacific. From seismic observatories throughout
the region, large earthquakes are reported to the Tsunami
Warning Center in Honolulu. Scientists at the center use tidal
gauges to determine whether a tsunami has formed. Within
an hour, a warning is issued. Although tsunami travel very
rapidly, there is sufficient time to evacuate all but the region
nearest the epicenter. For example, a tsunami generated near
the Aleutian Islands would take five hours to reach Hawaii,
and one generated near the coast of Chile would travel 15
hours before reaching Hawaii (Figure 8.21).
Students Sometimes Ask . . .
What is the largest wave triggered by an earthquake?
The largest wave ever recorded occurred in Lituya Bay, about
200 kilometers (125 miles) west of Juneau, Alaska’s capital. On
July 9, 1958, an earthquake triggered an enormous rockslide that
dumped 90 million tons of rock into the upper part of the bay. The
rockslide created a huge splash wave (different than a tsunami,
these waves are produced when an object splashes into water)
that swept over the ridge facing the rockslide and uprooted or
snapped off trees 1,740 feet above the bay. Even larger splash
waves have occurred in prehistoric times, including an estimated
3,000-foot wave that is thought to have resulted from a meteorite
impact in the Gulf of Mexico about 65 million years ago.
Landslides and Ground Subsidence
In the 1964 Alaskan earthquake, the greatest damage to structures was from landslides and ground subsidence triggered
by the vibrations, not from the vibrations themselves. At the
seaports of Valdez and Seward, the violent shaking caused
water logged sediments to experience liquefaction; the subsequent slumping carried both waterfronts into the sea. Because the disaster could happen again, the entire town of
Valdez was relocated about 7 kilometers (4 miles) away on
more stable ground. In Valdez, 31 people on a dock died
when it was jettisoned into the sea.
Most of the damage in Anchorage was attributed to landslides caused by shaking and lurching ground. Many homes
were destroyed in Turnagain Heights when a layer of clay
lost its strength and more than 200 acres of land slid toward
the ocean (Figure 8.22). A portion of this landslide has been
left undisturbed to serve as a reminder of this destructive
event. The site was named “Earthquake Park.” Downtown
Anchorage was also disrupted as sections of the main business district dropped by as much as 3 meters (10 feet).
FIGURE 8.22 Turnagain Heights slide caused by the 1964 Alaskan
earthquake. A. Vibrations from the earthquake caused cracks to appear
near the edge of the bluff. B. Within seconds blocks of land began to
slide toward the sea on a weak layer of clay. In less than 5 minutes much
of the Turnagain Heights bluff area had been destroyed. C. Photo of a
small portion of the Turnagain Heights slide. (Photo courtesy of U.S.
Geological Survey)
Turnagain Heights
FIGURE 8.21 Tsunami travel times to Honolulu, Hawaii, from selected
locations throughout the Pacific. (Data from NOAA)
200 meters
Original
profile
Asia
Aleutian Is.
North
America
9
6
7
5 hr
4 hr
3 h hr
2h r
1 hr r
10
8 hr
h
hr r
3 hr
hr
B.
12
11 hr
10 hr
hr
9
h
8 r
hr
7
h
6 r
hr
5h
4h r
3 r
h
2 r
h
1 r
hr
2 hr
4 hr
5 hr
Honolulu
New
Guinea
Australia
Layers of sand
Bootlegger
and gravel
Cove clay
Cracks developed
A.
1 hr
235
South
America
C.
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Earthquakes and Earth’s Interior
Fire
The 1906 San Francisco earthquake reminds us of the formidable threat of fire. The central city contained mostly large, older
wooden structures and brick-clad buildings that were mostly
destroyed by fires that started when gas and electrical lines
were severed. The fires raged uncontrolled for three days and
devastated over 500 city blocks (see Figure 8.3). The problem
was compounded by the initial ground shaking, which broke
the city’s water lines into hundreds of unconnected pieces.
The fire was finally contained when buildings were dynamited along a wide boulevard to create a fire break, the same
strategy used in fighting a forest fire. Although only a few
deaths were attributed to the fires, such is not always the case.
A 1923 earthquake in Japan (their worst quake prior to the
1995 Kobe tremor) triggered an estimated 250 fires, which devastated the city of Yokohama and destroyed more than half
the homes in Tokyo. More than 100,000 deaths were attributed to the fires, which were driven by unusually high winds.
Can Earthquakes be Predicted?
The vibrations that shook Northridge, California, in 1994 inflicted 57 deaths and about $40 billion in damage (Figure
8.23). This was from a brief earthquake (about 40 seconds) of
moderate rating 1MW 6.72. Seismologists warn that earth-
quakes of comparable or greater strength will occur along the
San Andreas Fault, which cuts a 1,300-kilometer (800-mile)
path through the state. The obvious question is, can earthquakes be predicted?
Short-Range Predictions
The goal of short-range earthquake prediction is to provide
a warning of the location and magnitude of a large earthquake within a narrow time frame. Substantial efforts to
achieve this objective are being put forth in Japan, the United
States, China, and Russia—countries where earthquake risks
are high (Table 8.3). This research has concentrated on monitoring possible precursors—phenomena that precede and thus
provide a warning of a forthcoming earthquake. In California, for example, some seismologists are measuring uplift,
subsidence, and strain in the rocks near active faults. Some
Japanese scientists are studying peculiar anomalous behavior that may precede a quake.
One claim of a successful short-range prediction was made
by Chinese seismologists after the February 4, 1975, earthquake in Liaoning Province. According to reports, very few
people were killed, although more than 1 million lived near
the epicenter, because the earthquake was predicted and the
population was evacuated. Recently, some Western seismologists have questioned this claim and suggest instead that an
FIGURE 8.23 Damage to Interstate 5 caused by the January 17, 1994, Northridge earthquake.
(Photo by Tom McHugh/Photo Researchers, Inc.)
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Can Earthquakes Be Predicted?
TABLE 8.3
Some Notable Earthquakes
Year
1556
1755
*1811–1812
*1886
*1906
1908
1923
1960
*1964
1970
*1971
1975
1976
1985
1988
*1989
1990
1993
*1994
1995
1999
1999
2001
2003
2004
2005
237
Location
Shensi, China
Lisbon, Portugal
New Madrid, Missouri
Charleston, South Carolina
San Francisco, California
Messina, Italy
Tokyo, Japan
Southern Chile
Alaska
Peru
San Fernando, California
Liaoning Province, China
Tangshan, China
Mexico City
Armenia
San Francisco Bay area
Iran
Latur, India
Northridge, California
Kobe, Japan
Izmit, Turkey
Chi-Chi, Taiwan
Bhuj, India
Bam, Iran
Indian Ocean
Pakistan/Kashmir
Deaths(est.)
Magnitude
830,000
70,000
Few
60
1,500
120,000
143,000
5,700
131
66,000
65
1,328
240,000
9,500
25,000
62
50,000
10,000
51
5,472
17,127
2,300
25,000+
41,000+
230,000
83,000
Unknown
Unknown
Unknown
Unknown
8.3
Unknown
7.9
9.5
9.2
7.8
6.5
7.5
7.6
8.1
6.9
7.1
7.3
6.4
6.7
6.9
7.4
7.6
7.9
6.6
9.0
7.6
Comments
Possibly the greatest natural disaster.
Tsunami damage extensive.
Three major earthquakes.
Greatest historical earthquake in the eastern United States.
Fires caused extensive damage.
Fire caused extensive destruction.
The largest-magnitude earthquake ever recorded.
Greatest North American earthquake.
Great rockslide.
Damage exceeded $1 billion.
First major earthquake to be predicted.
Not predicted.
Major damage occurred 400 km from epicenter.
Poor construction practices.
Damages exceeded $6 billion.
Landslides and poor construction practices caused great damage.
Located in stable continental interior.
Damages in excess of $15 billion.
Damages estimated to exceed $100 billion.
Nearly 44,000 injured and more than 250,000 displaced.
Severe destruction; 8,700 injuries.
Millions homeless.
Ancient city with poor construction.
Devastating tsunami damage.
Many landslides; 4 million homeless.
*U.S. earthquakes.
Source: U.S. National Oceanic and Atmospheric Administration
intense swarm of foreshocks that began 24 hours before the
main earthquake may have caused many people to evacuate
spontaneously. Furthermore, an official Chinese government
report issued 10 years later stated that 1,328 people died and
16,980 injuries resulted from this earthquake.
One year after the Liaoning earthquake at least 240,000
people died in the Tangshan, China, earthquake, which was
not predicted. The Chinese have also issued false alarms. In
a province near Hong Kong, people reportedly left their
dwellings for over a month, but no earthquake followed.
Clearly, whatever method the Chinese employ for short-range
predictions, it is not reliable.
For a short-range prediction scheme to warrant general
acceptance, it must be both accurate and reliable. Thus, it must
have a small range of uncertainty as regards to location and timing,
and it must produce few failures, or false alarms. Can you imagine the debate that would precede an order to evacuate a large
city in the United States, such as Los Angeles or San Francisco? The cost of evacuating millions of people, arranging
for living accommodations, and providing for their lost work
time and wages would be staggering.
Long-Range Forecasts
In contrast to short-range predictions, which aim to predict
earthquakes within a time frame of hours or at most days,
long-range forecasts give the probability of a certain magni-
tude earthquake occurring on a time scale of 30 to 100 years
or more. Stated another way, these forecasts give statistical
estimates of the expected intensity of ground motion for a
given area over a specified time frame. Although long-range
forecasts may not be as informative as we might like, the data
are important for updating the Uniform Building Code,
which contains nationwide standards for designing earthquake-resistant structures.
Long-range forecasts are based on the premise that earthquakes are repetitive or cyclical, like the weather. In other
words, as soon as one earthquake is over, the continuing motions of Earth’s plates begin to build strain in the rocks again,
until they fail once more. This has led seismologists to study
historical records of earthquakes to see if there are any discernible patterns so that the probability of recurrence might
be established.
One study conducted by the U.S. Geological Survey gives
the probability of a rupture occurring along various segments
of the San Andreas Fault for the 30 years between 1988 and
2018 (Figure 8.24). From this investigation, the Santa Cruz
Mountains area was given a 30 percent probability of producing a 6.5-magnitude earthquake during this time period.
In fact, it produced the Loma Prieta quake in 1989, of 7.1
magnitude.
The region along the San Andreas Fault given the highest
probability (90 percent) of generating a quake is the Parkfield
section. This area has been called the “Old Faithful” of
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Earthquakes and Earth’s Interior
North
Coast
San
Francisco
Peninsula
Probability of a
major earthquake
from 1988 to 2018
on the San Andreas
Fault
South
Santa Cruz
Mountains
90%
Less
than
10%
Parkfield
Cholame
Mojave
Carrizo
20%
30%
San Andreas
Fault
40%
20%
Segment of
very low
probability
10%
30%
Epicenter of
Oct. 17, 1989,
earthquake
Coachella
Valley
30%
San Francisco
San
Bernardino
Mountains
There are also variations in composition and temperature with depth that indicate the interior of our planet is very
dynamic. The rocks of the mantle and
crust are in constant motion, not only
moving about through plate tectonics, but
also continuously recycling between the
surface and the deep interior. Furthermore, it is from Earth’s deep interior that
the water and air of our oceans and atmosphere are replenished, allowing life to
exist at the surface.
Formation of Earth’s Layered
Structure
As material accumulated to form Earth
(and for a short period afterward), the
Los Angeles
high-velocity impact of nebular debris and
the decay of radioactive elements caused
FIGURE 8.24 Probability of a major earthquake from 1988 to 2018 along the San Andreas
the temperature of our planet to steadily
Fault.
increase. During this time of intense heating, Earth became hot enough that iron and
nickel began to melt. Melting produced liquid blobs of heavy
earthquake zones because activity here has been very regular
metal that sank toward the center of the planet. This process
since record keeping began in 1857. In late September 2004, a
occurred rapidly on the scale of geologic time and produced
magnitude 6.0 earthquake again struck this area. Although
Earth’s dense iron-rich core.
the event was more than a decade overdue, it did demonstrate
The early period of heating resulted in another process of
the potential usefulness of long-range forcasts. Another section
chemical differentiation, whereby melting formed buoyant
between Parkfield and the Santa Cruz Mountains is given a
masses of molten rock that rose toward the surface, where
very low probability of generating an earthquake. This area
they solidified to produce a primitive crust. These rocky mahas experienced very little seismic activity in historical times;
terials were rich in oxygen and “oxygen-seeking” elements,
rather, it exhibits a slow, continual movement known as fault
particularly silicon and aluminum, along with lesser amounts
creep. Such movement is beneficial because it prevents strain
of calcium, sodium, potassium, iron, and magnesium. In adfrom building to high levels in the rocks.
dition, some heavy metals such as gold, lead, and uranium,
In summary, it appears that the best prospects for making
which have low melting points or were highly soluble in the
useful earthquake predictions involve forecasting magnitudes
ascending molten masses, were scavenged from Earth’s inand locations in time scales of years, or perhaps even decades.
terior and concentrated in the developing crust. This early
These forecasts are important because they provide inforperiod of chemical segregation established the three basic dimation used to develop the Uniform Building Code and asvisions of Earth’s interior—the iron-rich core; the thin primitive
sist in land-use planning.
crust; and Earth’s largest layer, called the mantle, which is located between the core and crust (Figure 8.25).
Earth’s Interior
Forces Within
Earth’s Interior
If you could slice Earth in half, the first thing you would notice is that it has three distinct layers. The heaviest materials
(metals) would be in the center. Lighter solids (rocks) would
be in the middle and liquids and gases would be on top.
Within Earth we know these layers as the iron core, the rocky
mantle and crust, the liquid ocean, and the gaseous atmosphere. More than 95 percent of the variations in composition and temperature within Earth are due to layering.
However, this is not the end of the story. If it were, Earth
would be a dead, lifeless cinder floating in space.
Earth’s Internal Structure
In addition to these three compositionally distinct layers,
Earth can be divided into layers based on physical properties. The physical properties used to define such zones include whether the layer is solid or liquid and how weak or
strong it is. Knowledge of both types of layers is essential to
our understanding of basic geologic processes, such as volcanism, earthquakes, and mountain building (Figure 8.25).
Earth’s Crust The crust, Earth’s relatively thin, rocky outer
skin, is of two types: continental crust and oceanic crust. Both
share the word crust, but the similarity ends there. The
oceanic crust is roughly 7 kilometers (4 miles) thick and
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Earth’s Interior
Hydrosphere
(liquid)
Atmosphere
(gas)
Layering by
Physical Properties
Layering by
Chemical Composition
Oceanic
crust
Continental
crust
Crust
(low density rock
7–70 km thick)
Lithosphere
Up
pe
rm
an
tl e
(s
Lithosphere
(solid and brittle
100 km thick)
Asthenosphere
(solid, but mobile)
id
ol
)
Lower mantle
(solid)
Mantle
(high density rock)
0k
m
90
28
Upper
mantle
m
Inner
core
(solid)
515
0
660 km
k
289
Outer core
(liquid)
km
410 km
660 km
239
Core
(iron + nickel)
63
71
km
FIGURE 8.25 Views of Earth’s layered structure. The right side of the large cross section shows
that Earth’s interior is divided into three different layers based on compositional differences—the
crust, mantle, and core. The left side of the large cross section depicts the five main layers of
Earth’s interior based on physical properties and hence mechanical strength—the lithosphere,
asthenosphere, lower mantle, outer core, and inner core. The block diagram to the left of the large
cross section shows enlarged views of the upper portion of Earth’s interior.
composed of the dark igneous rock basalt. By contrast, the
continental crust averages 35–40 kilometers (22–25 miles)
thick but may exceed 70 kilometers (40 miles) in some mountainous regions such as the Rockies and Himalayas. Unlike
the oceanic crust, which has a relatively homogeneous chemical composition, the continental crust consists of many rock
types. Although the upper crust has an average composition
of a granitic rock called granodiorite, it varies considerably from
place to place.
Continental rocks have an average density of about 2.7
grams per cubic centimeter, and some are 4 billion years old.
The rocks of the oceanic crust are younger (180 million years
or less) and denser (about 3.0 grams per cubic centimeter)
than continental rocks.*
*Liquid water has a density of 1 gram per cubic centimeter; therefore, the density of
basalt is three times that of water.
Earth’s Mantle More than 82 percent of Earth’s volume is
contained in the mantle, a solid, rocky shell that extends to a
depth of about 2,900 kilometers (1,800 miles). The boundary
between the crust and mantle represents a marked change in
chemical composition. The dominant rock type in the uppermost mantle is peridotite, which is richer in the metals magnesium and iron than the minerals found in either the
continental or oceanic crust.
The upper mantle extends from the crust-mantle boundary
down to a depth of about 660 kilometers (410 miles). The
upper mantle can be divided into two different parts. The top
portion of the upper mantle is part of the stiff lithosphere, and
beneath that is the weaker asthenosphere.
The lithosphere (sphere of rock) consists of the entire crust
and uppermost mantle and forms Earth’s relatively cool, rigid
outer shell. Averaging about 100 kilometers (62 miles) in
thickness, the lithosphere is more than 250 kilometers (155
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miles) thick below the oldest portions of the continents
(Figure 8.25). Beneath this stiff layer to a depth of about 350
kilometers (217 miles) lies a soft, comparatively weak layer
known as the asthenosphere (“weak sphere”). The top portion of the asthenosphere has a temperature/pressure regime
that results in a small amount of melting. Within this very
weak zone, the lithosphere is mechanically detached from
the layer below. The result is that the lithosphere is able to
move independently of the asthenosphere.
It is important to emphasize that the strength of various
Earth materials is a function of both their composition and
the temperature and pressure of their environment. The entire lithosphere does not behave like a brittle solid similar to
rocks found on the surface. Rather, the rocks of the lithosphere
get progressively hotter and weaker (more easily deformed)
with increasing depth. At the depth of the uppermost asthenosphere, the rocks are close enough to their melting temperature that they are very easily deformed, and some
melting may actually occur. Thus, the uppermost asthenosphere is weak because it is near its melting point, just as hot
wax is weaker than cold wax.
From 660 kilometers (410 miles) deep to the top of the core,
at a depth of 2,900 kilometers (1,800 miles), is the lower mantle. Because of an increase in pressure (caused by the weight
of the rock above), the mantle gradually strengthens with
depth. Despite their strength, however, the rocks within the
lower mantle are very hot and capable of very gradual flow.
Earth’s Core The composition of the core is thought to be an
iron-nickel alloy with minor amounts of oxygen, silicon, and
sulfur—elements that readily form compounds with iron. At
the extreme pressure found in the core, this iron-rich material
has an average density of nearly 11 grams per cubic centimeter and approaches 14 times the density of water at Earth’s
center.
The core is divided into two regions that exhibit very different mechanical strengths. The outer core is a liquid layer
2,270 kilometers (1,410 miles) thick. It is the movement of
metallic iron within this zone that generates Earth’s magnetic
field. The inner core is a sphere with a radius of 1,216 kilometers (754 miles). Despite its higher temperature, the iron in
the inner core is solid due to the immense pressures that exist
in the center of the planet.
Surprisingly, meteorites—known as shooting stars—that
collide with Earth provide evidence of Earth’s inner composition. Because meteorites are part of the solar system, they
are assumed to be representative samples. Their composition
ranges from metallic meteorites made of iron and nickel to
stony meteorites composed of dense silicate minerals similar to those found in the rock peridotite. Because the crust
contains a much smaller percentage of iron than do meteorites, geologists believe that most of Earth’s iron and other
dense metals sank toward the center of the planet in its early
history. The surrounding mantle is composed of rock similar
to stony meteorites.
Strong evidence for a molten iron outer core is provided by
Earth’s magnetic field. The most widely accepted mechanism
explaining Earth’s magnetic field requires that a large por-
tion of the core be made of a material that conducts electricity, such as iron, and further that it is mobile enough to flow.
Both of these conditions are met by our current model of
Earth’s core—a solid inner core and mobile liquid outer core.
Probing Earth’s Interior:
“Seeing” Seismic Waves
Discovering the structure and properties of Earth’s deep interior has not been easy (see Box 8.2). Light does not travel
through rock, so we must find other ways to “see” into our
planet.
The best way to learn about Earth’s interior is to dig or
drill a hole and examine it directly. Unfortunately, this is only
possible at shallow depths. The deepest a drilling rig has ever
penetrated is only 12.3 kilometers (8 miles), which is about
1/500 of the way to Earth’s center! Even this was an extraordinary accomplishment because temperature and pressure
increase so rapidly with depth.
Fortunately, many earthquakes are large enough that their
seismic waves travel all the way through Earth and can be
recorded on the other side (Figure 8.26). This means that the
seismic waves act like medical x-rays used to take images of
a person’s insides. There are about 100 to 200 earthquakes
each year that are large enough (about M w 7 6) to be well
recorded by seismographs all around the globe. These large
earthquakes provide the means to “see” into our planet and
have been the source of most of the data that allowed us to
figure out the nature of Earth’s interior.
FIGURE 8.26 Slice through Earth’s interior showing some of the ray
paths that seismic waves from an earthquake would take. Notice that in
the mantle, the rays follow curved (refracting) paths rather than straight
paths because the seismic velocity of rocks increases with depth, a result
of increasing pressure with depth.
Earthquake
Inner
core
Outer
core
Mantle
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241
BOX 8.2 UNDERSTANDING EARTH
Recreating the Deep Earth
Seismology alone cannot determine what
Earth is made of. Additional information
must be obtained by some other means so
that the seismic velocities can be interpreted
in terms of rock type. This is done using
mineral physics experiments performed in
laboratories. By squeezing and heating minerals and rocks, physical properties like stiffness, compressibility, and density (and
therefore seismic velocities) can be directly
measured. This means that the conditions of
the mantle and core can be simulated, and
the results compared to seismic modeling.
Most mineral physics experiments are
done using giant presses involving very
hard carbonized steel. The highest pressures, however, are obtained using diamond-anvil presses like the one shown in
Figure 8.B. These take advantage of two important characteristics of diamonds—their
hardness and transparency. The tips of two
diamonds are cut off, and a small sample of
mineral or rock is placed in between. Pressures as great as those in the interior of
Jupiter have been obtained by squeezing the
two diamonds together. High temperatures
are achieved by firing a laser beam through
the diamond and into the mineral sample.
Besides measuring seismic velocities at
the conditions of different depths within
Earth, there are other important mineral
FIGURE 8.B High-pressure experiments inside a diamond anvil cell (left photo) can recreate the
conditions at the center of a planet. The whole apparatus is small and can fit on top of a table.
High pressures are generated by cutting the tips off high-quality diamonds (right photo), putting
a small sample of rock between, squeezing the diamonds together and heating the sample
with a laser. (Left photo courtesy of Lawrence Livermore National Laboratory; right photo
by Douglass L. Peck Photography)
physics experiments. One experiment determines the temperature at which minerals will begin to melt under various
pressures. Another experiment determines
(at different temperatures) the pressures at
which one mineral phase will become un-
Interpreting the waves recorded on seismograms in order
to identify the structure of Earth’s interior is complicated by
the fact that seismic waves usually do not travel along straight
paths. Instead, seismic waves are reflected, refracted, and diffracted as they pass through our planet. They reflect off
boundaries between different layers, they refract (or bend)
when passing from one layer to another layer, and they diffract around any obstacles they encounter (Figure 8.26). These
different wave behaviors have been used to identify the
boundaries that exist within Earth.
One of the most noticeable behaviors of seismic waves is
that they follow strongly curved paths (Figure 8.26). This occurs because the velocity of seismic waves generally increases
with depth. In addition, seismic waves travel faster when rock
is stiffer or less compressible. These properties of stiffness and
compressibility are then used to interpret the composition and
temperature of the rock. For instance, when rock is hotter, it becomes less stiff (imagine taking a frozen chocolate bar and
then heating it up!), and waves travel more slowly. Waves also
travel at different speeds through rocks of different compositions. Thus, the speed that seismic waves travel can help determine both the kind of rock that is inside Earth and how hot
it is.
stable and convert into a new high-pressure
phase. Yet another involves making these
same tests for slightly different mineral
compositions. All of these experiments are
needed because there are changes in composition and temperature within Earth.
Discovering Boundaries: The Moho
The boundary between the crust and mantle, called the
Moho, was one of the first features of Earth’s interior discovered using seismic waves. Croatian seismologist Andrija
Mohoroviĉiĉ discovered this boundary in 1909 and it is
named in his honor. At the base of the continents P waves
travel about 6 km/s but abruptly increase to 8 km/s at a
slightly greater depth.
Mohoroviĉiĉ cleverly used this jump in seismic velocity
to identify the boundary. He noticed that there were two different sets of seismic waves that were recorded within a few
hundred kilometers of an earthquake. One set of waves
moved across the ground at about 6 km/s, and the other set
of waves moved across the ground at about 8 km/s. From
this data, Mohoroviĉiĉ correctly determined that the different
waves were coming from two different layers, as shown in
Figure 8.27.
When a shallow earthquake occurs, there is a direct wave
that moves straight through the crust and is recorded at
nearby seismographs. Seismic waves also follow a path down
through the crust and along the top of the mantle. These are
called refracted waves because they are bent, or refracted, as
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Seismograph
#1
Seismograph
#2
Seismograph
#3
Refracted
wave
Moho
A.
Earthquake
#1
Direct wave
#2
#3
Refracted
wave
Moho
B.
Earthquake
#1
#2
Direct wave
#3
Refracted
wave
Moho
C.
FIGURE 8.27 Diagram showing seismic waves from an earthquake arriving at three different
seismographs. Over a short distance, such as at seismograph #1, the direct wave arrives first. For
greater distances, such as at seismograph #3, the refracted wave arrives first. The distance at which
both waves arrive at the same time can be used to determine the thickness of the crust.
they enter the mantle. Refracted P waves travel across the
ground at the speed of the waves in the mantle (8 km/s). At
nearby locations the direct wave arrives first (Figure 8.27).
However, at greater distances the refracted wave is the first
to arrive. The distance at which both waves arrive at the same
time can be used to determine the depth of the Moho. Thus,
using just these two sets of waves and an array of seismographs, the thickness of the crust at any location can be
determined.
The difference in arrival times between direct and refracted
waves is analogous to driving along local roads versus taking the interstate. For short distances, you will arrive faster if
you take local roads. For greater distances, it takes less time
if you first drive to a major highway and then travel along it.
The point where both routes take the same amount of time is
directly related to how far you are from a major highway. Or,
when determining the depth of the Moho, how far the mantle (fast layer) is from the surface.
Discovering Boundaries:
The Core–Mantle Boundary
Evidence that Earth has a distinct central core was uncovered
in 1906 by a British geologist, Richard Dixon Oldham. (In
1914 Beno Gutenberg calculated the depth to the core bound-
ary as 2900 kilometers, a value that has
stood the test of time.) Oldham observed that at distances of more than
about 100° from a large earthquake P
and S waves were absent, or very
weak. In other words, the central core
produced a “shadow zone” for seismic
waves as shown in Figure 8.28.
As Oldham predicted, Earth’s core
exhibits markedly different elastic
properties from the mantle above,
which causes considerable refraction of
P waves—similar to how light is refracted (bent) as it passes from air to
water. In addition, because the outer
core is liquid iron, it blocks the transmission of S waves (recall S waves do
not travel through liquids).
Figure 8.28 shows the locations of
the P and S wave shadow zones and
how the paths of the waves are affected
by the core. While there are still P and
S waves that arrive in the shadow
zone, they are all very different than
would be expected for a planet without a core.
Discovering Boundaries:
The Inner Core–Outer Core
Boundary
The boundary between the solid inner
core and liquid outer core was discovered in 1936 by Danish
seismologist Inge Lehman. She could not tell whether the
inner core was actually solid or not, but using basic trigonometry reasoned that some P waves were being strongly refracted by a sudden increase in seismic velocities at the inner
core–outer core boundary. This is the opposite situation as
Students Sometimes Ask . . .
As compared to continental crust, oceanic crust is quite thin.
Has there ever been an attempt to drill through it to obtain a
sample of the mantle?
Yes. Project Mohole was initiated in 1958 to retrieve a sample of
material from Earth’s mantle by drilling a hole through Earth’s
crust to the Mohoroviĉiĉ discontinuity, or Moho. The plan was to
drill to the Moho to gain valuable information on Earth’s age,
makeup, and internal processes. Despite a successful test phase,
drilling was halted as control of the project was shifted from one
governmental organization to another. Although Project Mohole
failed in its intended purpose, it did show that deep-ocean
drilling is a viable means of obtaining geological samples.
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Chapter Summary
Earthquake
Earthquake
Direct
P waves
Direct
S waves
Inner
core
Outer
core
Outer
core
Inner
core
Mantle
100°
100°
100°
100°
Mantle
P-wave
shadow
zone
140°
140°
180°
P waves
recorded
A. P-wave shadow zone
S-wave shadow zone
(weak diffracted waves)
B. S-wave shadow zone
FIGURE 8.28 Two views of Earth’s interior showing the effects of the outer and inner cores on the ray
paths of P and S waves. A. When P waves interact with the slow-velocity liquid iron of the outer core,
their rays are refracted downward. This creates a shadow zone where no direct P waves are recorded
(although diffracted P waves travel there). The P waves that travel through the core are called PKP
waves. The “K” represents the path through the core, and comes from the German word for core, which
is kern. The increase in seismic velocity at the top of the inner core can refract waves sharply so that
some arrive within the shadow zone, which is shown here as a single ray. B. The core is an obstacle to
S waves, because they cannot pass through liquids. Therefore, a large shadow zone exists for S waves.
However, some S waves diffract around the core and are recorded on the other side of the planet.
what occurs to produce the P-wave shadow zone. When seismic velocities suddenly decrease, such as at the mantle–outer
core boundary, waves get bent so that there is a shadow zone
where no direct waves arrive. When seismic waves suddenly
increase, as they do at the outer core–inner core boundary,
waves get bent so that several P waves can sometimes arrive
at a single location. In the case of the inner core, these waves
can even be refracted enough to arrive within the P-wave
shadow zone. Both of these occurrences, shown in Figure
8.28A, are proof of a distinct inner core.
In summary, although earthquakes can be very destructive, much of our knowledge about Earth’s interior comes
from the study of these phenomena. As our knowledge of
earthquakes and their causes improve, we learn more of our
planet’s internal workings and possibly how to reduce the
consequences of tremors.
Chapter Summary
Earthquakes are vibrations of Earth produced by the rapid
release of energy from rocks that rupture because they
have been subjected to stresses beyond their limit. This
energy, which takes the form of waves, radiates in all directions from the earthquake’s source, called the focus. The
movements that produce most earthquakes occur along
large fractures, called faults, that are associated with plate
boundaries.
Two main groups of seismic waves are generated during
an earthquake: (1) surface waves, which travel along the
outer layer of Earth; and (2) body waves, which travel
through Earth’s interior. Body waves are further divided
into primary, or P, waves, which push (compress) and pull
(expand) rocks in the direction the wave is traveling, and
secondary, or S, waves, which shake the particles in rock at
right angles to their direction of travel. P waves can travel
through solids, liquids, and gases. Fluids (gases and liquids) will not transmit S waves. In any solid material,
P waves travel about 1.7 times faster than do S waves.
The location on Earth’s surface directly above the focus
of an earthquake is the epicenter. An epicenter is determined using the difference in velocities of P and S waves.
There is a close correlation between earthquake epicenters and
plate boundaries. The principal earthquake epicenter zones
are along the outer margin of the Pacific Ocean, known
as the circum-Pacific belt and through the world’s oceans
along the oceanic ridge system.
Seismologists use two fundamentally different measures to
describe the size of an earthquake: intensity and magnitude. Intensity is a measure of the degree of ground
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shaking at a given locale based on the amount of damage.
The Modified Mercalli Intensity Scale uses damages to buildings in California to estimate the intensity of ground shaking for a local earthquake. Magnitude is calculated from
seismic records and estimates the amount of energy released at the source of an earthquake. Using the Richter
scale, the magnitude of an earthquake is estimated by measuring the amplitude (maximum displacement) of the largest
seismic wave recorded. A logarithmic scale is used to express magnitude, in which a tenfold increase in ground
shaking corresponds to an increase of 1 on the magnitude
scale. Moment magnitude is currently used to estimate the
size of moderate and large earthquakes. It is calculated
using the average displacement of the fault, the area of the
fault surface, and the sheer strength of the faulted rock.
The most obvious factors that determine the amount of
destruction accompanying an earthquake are the magnitude of the earthquake and the proximity of the quake to
a populated area. Structural damage attributable to earthquake vibrations depends on several factors, including
(1) intensity, (2) duration of the vibrations, (3) nature of the
material upon which the structure rests, and (4) the design
of the structure. Secondary effects of earthquakes include
tsunamis, landslides, ground subsidence, and fire.
Substantial research to predict earthquakes is under way
in Japan, the United States, China, and Russia—countries
where earthquake risk is high. No consistent method of
short-range prediction has yet been devised. Long-range
forecasts are based on the premise that earthquakes are
repetitive or cyclical. Seismologists study the history of
earthquakes for patterns, so their occurrences might be
predicted.
Earth’s internal structure is divided into layers based on
differences in chemical composition and on the basis of
changes in physical properties. Compositionally, Earth is
divided into a thin outer crust, a solid rocky mantle, and a
dense core. Based on physical properties, the layers of
Earth are (1) the lithosphere—the cool, rigid, outermost
layer that averages about 100 kilometers thick; (2) the
asthenosphere, a relatively weak layer located in the mantle beneath the lithosphere; (3) the more rigid lower mantle where rocks are very hot and capable of very gradual
flow; (4) the liquid outer core, where Earth’s magentic field
is generated; and (5) the solid inner core.
The continental crust has an average composition of a
granitic rock called granodiorite, whereas the oceanic crust
has a basaltic composition. Rocks similar to peridotite make
up the mantle. The core is made up mainly of iron and
nickel. An iron core explains the high density of Earth’s
interior as well as Earth’s magnetic field.
Key Terms
aftershock (p. 223)
asthenosphere (p. 240)
body wave (p. 225)
core (p. 240)
crust (p. 238)
earthquake (p. 220)
elastic rebound (p. 223)
epicenter (p. 226)
fault (p. 221)
fault creep (p. 224)
focus (p. 220)
foreshock (p. 223)
inner core (p. 240)
intensity (p. 228)
liquefaction (p. 232)
lithosphere (p. 239)
lower mantle (p. 240)
magnitude (p. 228)
mantle (p. 239)
Modified Mercalli Intensity
Scale (p. 228)
Mohoroviĉiĉ discontinuity
(Moho) (p. 241)
moment magnitude (p. 230)
outer core (p. 240)
primary (P) wave (p. 225)
Richter scale (p. 229)
secondary (S) wave (p. 225)
seismic sea wave (tsunami)
(p. 232)
seismogram (p. 225)
seismograph (p. 225)
seismology (p. 225)
surface wave (p. 225)
Review Questions
1. What is an earthquake? Under what circumstances do
earthquakes occur?
2. How are faults, foci, and epicenters related?
3. Who was first to explain the actual mechanism by which
earthquakes are generated?
4. Explain what is meant by elastic rebound.
5. Faults that are experiencing no active creep may be considered safe. Rebut or defend this statement.
6. Describe the principle of a seismograph.
7. Using Figure 8.10B, determine the distance between an
earthquake and a seismic station if the first S wave arrives 3 minutes after the first P wave.
8. List the major differences between P and S waves.
9. Which type of seismic wave causes the greatest destruction to buildings?
10. Most strong earthquakes occur in a zone on the globe
known as the _____.
11. In 1988 an earthquake in Armenia caused approximately
25,000 deaths. A year later an earthquake in the Bay Area
of California, having the same Richter magnitude of 6.9,
had only 62 deaths. Suggest a reason for the greater loss
of life in the Armenian event.
12. An earthquake measuring 7 on the Richter scale releases
about _____ times more energy than an earthquake with
a magnitude of 6.
13. List three reasons why the moment magnitude scale has
gained wide acceptance among seismologists.
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14. List four factors that affect the amount of destruction
caused by seismic vibrations.
15. In addition to the destruction created directly by seismic
vibrations, list three other types of destruction associated
with earthquakes.
16. Distinguish between the Mercalli scale and the Richter
scale.
17. What is a tsunami? In what two ways do earthquakes
generate tsunami?
18. Cite some reasons why an earthquake with a moderate
magnitude might cause more extensive damage than a
quake with a high magnitude.
245
19. What evidence do we have that Earth’s outer core is
molten?
20. Contrast the physical makeup of the asthenosphere and
the lithosphere.
21. Why are meteorites considered important clues to the
composition of Earth’s interior?
22. Describe the composition of the following:
a.
b.
c.
d.
continental crust
oceanic crust
mantle
core
Examining the Earth System
What potentially disastrous phenomenon often occurs when
the energy of an earthquake is tranferred from the solid Earth
to the hydrosphere (ocean) at their interface on the floor of the
ocean? When the energy from this event is expended along a
coast, how might coastal lands and the biosphere be altered?
Online Study Guide
The Earth Science Website uses the resources and flexibility
of the Internet to aid in your study of the topics in this chapter. Written and developed by Earth science instructors, this
site will help improve your understanding of Earth science.
Visit http://www.prenhall.com/tarbuck and click on the cover
of Earth Science 12e to find:
• Online review quizzes
• Critical thinking exercises
• Links to chapter-specific Web resources
• Internet-wide key term searches
GEODe: Earth Science
GEODe: Earth Science makes studying more effective by reinforcing key concepts using animation, video, narration, in-
teractive exercises, and practice quizzes. A copy is included
with every copy of Earth Science 12e