Download Chapter 30: The Interior of the Earth

30. The Interior of the Earth
Sometime during your childhood you may have
begun digging a hole in the ground with the intention of
digging through to the other side of the earth. Sooner or
later (probably sooner) you became bored with the project and turned to other amusements, never aware of
how utterly preposterous your undertaking had been in
the first place. Even the deepest mines and drill holes
penetrate only a fraction of one percent of the thickness
of the earth. Exposed rocks that originated “deep” within the crust of the planet were really formed only 20 or
25 kilometers down—not much compared to the 6400
kilometer depth to the center! Occasionally rocks that
may have come from as deep as 200 kilometers are carried upward by ascending magmas and erupted onto the
surface. Although such samples of the interior do give
us important information, they are neither distributed
evenly enough over the surface nor are they nearly
numerous enough to provide us with a complete picture.
Our information must come mostly from instruments
that probe the interior without actually going there—
indirect evidence.
As two large lead spheres are brought near the small
spheres, the gravitational attraction between the small
and large spheres twists the wire. Because the force
required to twist the wire any given amount can be
determined experimentally, the force is known, as are
the two masses and the distance between attracting
spheres. All that remains is to solve for G in the equation F ! GmM/d2.
Light Source
Mirror
Gold
The Density of the Earth
Lead
In Chapter 10 density was defined as mass divided
by volume. If we could determine the density of the
earth, that might give us some clues about its internal
structure. It appears that the volume should be easy
enough to find because we know the earth’s size and
shape, but finding the mass looks more difficult. The
way it is done is a bit indirect, but it has been possible
for nearly two centuries.
The method requires first that the universal gravitational constant, G, be found. Recall that this is the constant in the equation for gravitational force (Chapter 4).
Sir Henry Cavendish (English physicist and chemist,
1731-1810) used what is now called a Cavendish balance for determining relative densities, and while he
probably did not actually determine G, the Cavendish
balance can be used for this purpose. It consists (Fig.
30.1) of two small gold or platinum spheres mounted on
opposite ends of a lightweight metal bar, which is suspended by a very fine wire. The wire has a mirror
mounted on it, and a light reflected from the mirror onto
a scale reveals even the smallest twisting of the wire.
Lead
Gold
Figure 30.1. The Cavendish balance, which can be used
for determining the universal gravitational constant and,
from that, the mass of the earth.
With G now determined, we can calculate the mass
of the earth. The force exerted by the earth on a onegram mass can be measured quite accurately, as can the
distance between the mass and the center of the earth.
Now knowing F, G, m, and d, we can solve for M, the
mass of the earth, and then divide by the volume to get
density, which turns out to be 5.5 grams per cubic centimeter (g/cm3). The range in the densities of common
rocks at the surface of the earth is about 2.5 g/cm3 to 3.5
g/cm3, with an average of about 2.7 g/cm3. If the mean
density of the planet is higher than the density at the surface, then density must increase with depth. Since we
know that pressure increases with depth (because of the
increasing weight above), it comes as no surprise that
density does also.
Figure 30.2 shows two possible models that are
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consistent with this observation. Either the density
could increase relatively continuously with depth, or the
earth could consist of discrete layers whose densities
increase with depth discontinuously. Our determination
of the average density does not discriminate between
the two models. (It may occur to you that a third model,
a layered one in which some of the deeper layers are
less dense than the more shallow ones, could be constructed to yield the proper average density. This, however, would create an unstable situation called a “density inversion,” in which the less dense layers underneath
would tend to flow slowly and rise upward over a long
period of time while the heavier layers above would
break up and sink.)
(a)
(a)
(b)
(b)
Figure 30.2. Two models of the interior structure of the
earth that are consistent with the density measurements.
In model (a) density increases smoothly toward the center of the earth. In model (b) the earth consists of discrete layers, each denser than the ones above it.
(c)
Seismic Waves and the Structure of the Earth
Figure 30.3. (a) A road crosses an unstressed block of
lithosphere (the earth’s brittle outer shell). (b) Forces in
the earth bend the lithosphere. (c) When the lithosphere
can no longer accommodate the strain, it breaks,
rebounding elastically, along a fault. (Deformation is
exaggerated here for emphasis.)
Although the earth seems hard and brittle to us, it is
actually a giant elastic sphere (but only slightly
deformable compared to, say, a rubber ball). When segments of the earth’s lithosphere are subjected to stresses
(the nature of which we shall discuss later), they react by
bending—to a point. When they reach a condition of
strain at which they can bend no further, they rupture
abruptly and then rebound somewhat like a rubber band
that has just been pulled apart. The phenomenon is
called “elastic rebound,” and it is illustrated in Figure
30.3. The break along which the failure of the lithosphere occurs is called a fault. (Because of elasticity,
faults do not need to involve extremely large segments
of the lithosphere or to extend through its entire thickness.)
When rocks break abruptly, either at the surface or
(more typically) in the subsurface, shock waves are generated. These are seismic waves (seismos is Greek for
earthquake), and they occur in three essentially different
types. Recall that we identified three types of waves in
Chapter 13—compressional waves, shear waves, and surface waves. Seismic waves are of all three types, but for
historical reasons geologists call two of them by different
names. Early seismologists observed that three kinds of
waves, each with its own identifiable characteristics,
arrived at their instruments after an earthquake. The ones
that arrived first (compression) were called primary
waves, now shortened to P waves. The next to arrive
(shear) were the secondary waves, or S waves. The last
were the surface waves. While surface waves do the bulk
of the damage during an earthquake, they travel only near
the surface, and they therefore tell us little about the interior of the earth. It is important to understand that all three
types of seismic waves are produced simultaneously and
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travel in all directions outward from the hypocenter.
The hypocenter is the point in the subsurface
where the disturbance actually occurred. The location
of an earthquake referred to in a news account is the
point on the surface directly above the hypocenter and
is called the epicenter (see Fig. 30.4). The speeds of
the P waves and the S waves not only differ from each
other but they also change with the properties of the
rocks through which they travel. Essentially, the stiffer
and/or more dense the rocks, the faster the waves travel
through them, and that makes it possible to learn something about the interior of the earth by analyzing the
speeds of waves that pass through different parts of it.
direction parallel to the length of the recording cylinder,
everything moves with it except the inertial mass, which
remains stationary (Newton’s First Law, Chapter 3)
because there are no horizontal forces acting on it. The
result is that the pen creates a record of the movement on
the chart. In order to completely determine the nature of
ground-shaking, three seismometers must be used: two
oriented perpendicular to each other that record horizontal motion and one that records vertical motion. This latter seismometer, using a design analogous to our crude
model, would have the mass hung from a spring and the
pen and chart to the side of it.
Epicenter
t
ul
a
f
Hypocenter
Mass
Pen
Drum
Figure 30.4. Some seismologic terminology. The disturbance is caused when rocks move at a fault which,
because the earth is elastic, may not extend all the way
to the surface and be visible. The location of the disturbance (that is, the origin of the seismic waves) is the
hypocenter, sometimes called the focus. Directly above
that is the epicenter. The arrows represent the directions
of seismic waves and are perpendicular to the wave
fronts. Henceforth, we shall use arrows to represent the
waves.
Bedrock
Figure 30.5. A crude seismometer design that illustrates
the principles involved. Modern instruments operate on
the same principles, but are considerably more sophisticated.
As long as seismic waves stay within a medium
whose elastic properties do not vary in any direction,
they travel in straight lines. However, if seismic waves
travel through media with different elastic properties
(that is, from rocks of one elasticity to rocks of another,
or through rocks with elasticity that is gradually changing with depth), they experience refraction—a phenomenon we recall from our previous discussion of wave
behavior. The difference is suggested by Figure 30.6.
Early seismic investigations showed that waves arriving
at seismographs farther and farther from an earthquake
were progressively earlier and earlier than expected,
based on speeds calculated from waves that traveled
Seismic waves are detected by seismometers (literally, “earthquake measurers”), which display the effects
of earthquakes on charts called seismographs (“earthquake writers”). Figure 30.5 shows how a seismometer
works in principle, although modern instruments are
more sophisticated and do not look much like this. The
large mass is attached to the bedrock of the earth through
a wire hung from a frame. Underneath the mass is a
cylindrical drum bearing the recording chart, which is
also attached to bedrock. A pen on the mass can write on
the chart. When the earth shakes back and forth in a
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(a)
(a)
(b)
(b)
0°
Epicenter
103°
Shadow
zone for
P waves
143°
103°
Shadow zone
for S waves
143°
Figure 30.6. The behavior of seismic waves in hypothetical planets made of (a) material with constant elastic properties and (b) material that gradually and continuously becomes stiffer or more dense with depth.
Shadow
zone for
P waves
Figure 30.7. (a) The seismic shadow zone for P and S
waves for a hypothetical earthquake in Alaska is shaded. Inside the shaded ring, on the opposite side of the
planet from the epicenter, only P waves are received, so
this is a shadow zone for S waves only. (b) The interpretation of the seismic shadow zone. Dark arrows represent the paths of P waves, and light ones represent the
paths of S waves. The darker inner sphere must be liquid because S waves are not transmitted through it, and
it is called the core. Note that the solid inner core, discussed in the text, is not shown because the absence of
S waves does not require it, and the evidence for it is not
given on this diagram.
shorter distances. Because the waves that traveled farther must have traveled through greater depths than
those that traveled lesser distances (by either model in
Fig. 30.6), they must have traveled faster as they got
deeper. It follows that the rocks must have become
stiffer and/or more dense with depth and that the waves
therefore followed the curved sorts of paths shown in
Figure 30.6b. It turns out that the rocks also become
denser as they become stiffer.
However, a general increase in density is not all of
the story. Early in this century a Croatian seismologist
named Adrija Mohorovicic (1857-1936) compared the
travel times of seismic waves arriving at seismometer
stations less than 200 kilometers and more than 200
kilometers from earthquakes in eastern Europe. He dis-
covered that the data revealed a discontinuity in the
seismic wave velocities a few tens of kilometers below
the surface. A seismic discontinuity is a distinct and
abrupt change in the velocities of seismic waves (and
therefore in the stiffnesses and densities of the rocks) as
opposed to a smooth and gradual change. Mohorovicic
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few weak P waves arrive within the shadow zone by
reflection from a discontinuity within the liquid core,
and anomalously early arrivals of P waves traveling
through the center of the earth revealed that discontinuity to be a boundary between the liquid outer core and
a solid inner core, at a depth of about 5100 kilometers.
Figure 30.8 depicts what the seismic data have shown us
about the gross structure of the earth to this point in our
discussion. It is, indeed, a layered planet with a very thin
crust, a thick mantle (about 84 percent of the volume
altogether), and a core that is liquid on the outside and
solid in the center. We shall return to one further very
important seismic observation later on, but we now turn
to other indirect evidence about the earth’s interior to see
what can be inferred about its chemical composition.
Crust
Mantle
Liquid outer
core
Solid
inner
core
The Composition of the Crust
We live on the crust, so it is directly accessible and
we know quite a bit about its composition. In Chapter
28 we learned that continental and oceanic crust are
markedly different. The continental crust consists of a
wide variety of rock types—igneous, sedimentary, and
metamorphic rocks of all sorts—but is granitic on the
average. The rocks consist mostly of fewer than 20
common minerals, most of which are silicates (that is,
they contain the SiO44– molecular ion). The oceanic
crust also consists almost entirely of silicates, but with
a lower overall proportion of silicon. Contrasting with
the variety of rock types found in the continental crust,
the oceanic crust is virtually all basalt—a single specific type of rock. The density of the continental crust is
around 2.7 g/cm3 on the average; that of the oceanic
crust is about 3.0 g/cm3.
Figure 30.8. The layered structure of the earth, as
revealed by seismic data.
had discovered the base of the layer we live on and confirmed that the earth consists of at least two layers. The
upper one is called the crust, and the one underneath it
is called the mantle. The boundary between them is
named in his honor—the Mohorovicic discontinuity,
generally shortened (for obvious reasons) to the Moho.
The continental crust is fairly thick (30-40 kilometers
under stable platforms and as much as 60 kilometers
under mountain ranges) compared to the oceanic crust
(about 5 to 10 kilometers).
Near this time, Beno Gutenberg (GermanAmerican seismologist, 1889-1960) was puzzling over
different data. Observations of seismograph records for
strong earthquakes occurring long distances from seismometer stations showed an even more interesting
behavior: In a broad band 103° to 143° from any earthquake epicenter, no seismic waves were detected at all,
and beyond 143°, only P waves were recorded. Figure
30.7 depicts the observation and the interpretation. The
band of seismic nonresponse is called the shadow zone.
The P waves that would have emerged within it have
been refracted away by some very significant discontinuity, to emerge beyond 143° from the epicenter. The
complete absence of S waves beyond 103° from the epicenter indicates that the discontinuity is a boundary
between the solid mantle and a layer below, called the
outer core, that is liquid. (Remember that S waves are
shear waves, and shear waves can be propagated only in
a solid.)
As early as 1936, examination of seismograph
records for earthquakes around the world showed that a
The Composition of the Mantle
The lavas erupted by volcanoes vary considerably
in composition and come from a substantial range of
depths. The ones that come from the deepest locations
sometimes carry with them chunks that do not resemble
rocks normally found at the surface of the earth. They
consist of silicate minerals, but in proportions not found
in crustal rocks; chemically, there is less silicon and
more magnesium than found in the crust. These pieces
have come from the upper mantle and, while they might
have undergone some changes on the way up, they are
essentially samples of that layer. The rock type is called
peridotite.
Occasionally, large meteors are not entirely burned
up as they encounter the atmosphere of the earth, and
pieces of them land on the surface as meteorites. Most
meteorites generally belong to one of two broad classes: iron meteorites and stony meteorites. It is thought
that they are remains of one or more planet-size objects
that disintegrated early in the history of the solar system
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and now orbit the sun as swarms of debris. If this is correct, then they may represent samples of a planet somewhat like the earth. The stony meteorites turn out to
consist mostly of peridotite. Further, the stony meteorites account for 80 to 90 percent of the meteorites
seen to fall and collected thereafter (although the iron
meteorites, which stand out because of their unusual
appearance, account for most of the accidental finds in
which no meteor was seen to fall). This percentage is
about as expected if meteorites come from a planet with
a mantle five or six times as voluminous as its core, as
is the case with the earth.
Laboratory experiments done under very high pressures and temperatures show that the minerals composing peridotite cannot be stable at the depths of the lower
mantle; these minerals must undergo chemical reactions
that produce denser minerals at least at two different
depths. In fact, seismic evidence reveals two zones of
rapid change from lower to higher seismic velocities
between depths of 300 and 700 kilometers, and it is
thought that these represent the two sets of reactions
predicted by laboratory experiments. Below that, the
seismic velocities simply increase gradually to the base
of the mantle. The chemical composition of the lower
mantle is probably nearly the same as that of the upper
mantle, but the atoms have rearranged themselves by
these chemical reactions to form denser, more stable
compounds called dense oxides.
currents exist in the core? The answer is that the outer
core is liquid with temperatures that vary from top to
bottom, and therefore it must experience convection. If
the core were made of some molten metal, then the
metal atoms flowing as convection currents would constitute a moving electrical conductor, and a magnetic
field would result.
Remember that there are two major classes of
meteorites, the stony meteorites and the iron meteorites,
and that the stony meteorites provide evidence for the
peridotitic composition of the upper mantle. The iron
meteorites are thought to be the remains of the cores of
disrupted planet-size bodies, and they may therefore be
similar in composition to the core of the earth. They are
mostly iron, alloyed with lesser amounts of nickel.
Such a composition would, as a convecting liquid,
account for the magnetic field.
Because the average density of the earth is about
5.5 g/cm3, and we have already concluded what the densities of the crust and mantle are, the density of the core
must be a value that would provide the proper average
density of the earth. The density of liquid iron (“lightened” by a small amount of one or more elements of
lower atomic number) at the extreme pressures that
exist at the depth of the core would be 10 or 11 g/cm3,
and the density of the solid would be 12 to 14 g/cm3.
These turn out to be about the values needed to provide
the observed average density of the earth.
Taken together, all of this is strong circumstantial
evidence that the core of the earth consists mostly of
iron, molten in the outer region but solid in the center.
It might seem strange that the inner core, which is
deeper and thus hotter than the outer core, would be
solid; whereas the cooler outer core, made of essentially
the same thing, would be molten. There are two reasons
for this. First, the inner core is not only hotter, but also
under greater pressure. Recall from Chapter 10 that the
difference between a solid and a liquid was in the distances and forces between the atoms or molecules that
constituted them; in a solid the atoms were closer together, and there were bonding forces (Chapters 20 and 21)
that held them together. At the pressures of the inner
core, the metal atoms are forced close enough together
that they form a solid. The other reason addresses why
the outer core stays liquid. When a liquid solidifies, we
say that it “crystallizes.” In order to get water to crystallize and become ice, energy must be removed from
the water. The same is true of any other liquid, including the liquid iron of the outer core. The heat energy
removed is called the latent heat of crystallization. As
the heaviest atoms of the outer core (mostly iron atoms)
slowly fall toward the center of the earth and reach the
inner-outer core boundary, they crystallize, releasing the
latent heat of crystallization. That heat keeps the outer
core molten (and also produces the convection currents
that generate the magnetic field).
The Composition of the Core
No pieces of the core have ever been entrained in
volcanic lavas or otherwise reached the surface of the
earth, so it might seem as if we could only speculate
with great uncertainty about its composition. However,
there are some observations we can make that put constraints on what that composition is likely to be, and
there are other observations that suggest possible compositions. Among these are the magnetic field of the
earth, the compositions of iron meteorites, and the average density of the earth.
The earth possesses a magnetic field that behaves
somewhat as if the planet had a large permanent magnet
within its core. Inasmuch as the inner core is solid, it
may occur to you that it might be made of magnetized
iron (or of something else that can become magnetized),
but there is a good reason that this cannot be so. If we
were to heat a piece of magnetic iron to a high enough
temperature, it would lose its magnetism; that temperature is called the Curie temperature, and for iron it is
760 °C. The temperature of the core of the earth is
much higher than the Curie temperature for any substance, and so nothing in the core could be permanently
magnetized. However, an electrical current in a wire
also creates a magnetic field, so perhaps electrical currents could generate the earth’s field. How could such
296
Summary of the Chemical Layering of the Earth—
Differentiation
that is partially molten—perhaps only 1–10 percent liquid, but enough to make the rock plastic. (Here, the
word plastic means nonrigid, deformable, and capable
of flowing in response to pressure that is applied
through long periods of time. Think of some substance
like Silly Putty which, when rolled into a ball and left
on a table overnight, becomes flattened just by the force
of gravity. This substance is not what you would usually think of as a fluid, but it will flow by slow deformation—that is, it is plastic. The low velocity zone is certainly not like “silly putty” in any other way, but it is
plastic.)
Because the low velocity zone is a weak, soft layer
in the earth, it is called the asthenosphere (from the
Greek astheneia, meaning weak), and Figure 30.9
shows why it exists. The curve that shows how the temperature of the earth changes with depth tends to flatten
out, so that the temperature changes more slowly as
depth increases. At depths between about 70 and 250
kilometers below the surface, this curve crosses into the
shaded area, which represents the range of temperatures
and pressures at which solid peridotite begins to melt.
(Like most rocks, peridotite consists of more than one
mineral, each of which has its own melting point at any
given pressure, and that is why the boundary between
solid and liquid peridotite is a band, rather than a sharp
line.) When the temperature curve crosses below that
The earth is a differentiated planet, which means
that it is divided into layers of different chemical compositions. We have seen (in Chapter 28) that other planets appear to be differentiated, too. In the earth, there is
a core that consists mostly of iron (the inner and outer
core may have slightly different compositions but, to a
first approximation, they are the same), a mantle that
consists of dense oxides (lower mantle) and peridotite
(upper mantle), and a crust consisting in some places of
basalt (oceans) and in others of many rock types that
have the average composition of granite (continents).
Outside of the crust, we have the hydrosphere (the
water that covers much of the earth’s surface) and finally the atmosphere (consisting of gases).
Note that the densities of these various layers
change from densest at the center to least dense at the
outside of the earth. This is intuitively satisfying—we
would somehow like the “heaviest” material to be closest to the center of the planet, and it is. This is what is
meant by a differentiated planet—not just that it is layered, but that the densest layers are on the inside.
Moreover, the layers are chemically different from one
another. That differentiation should come about seems
logical, but just how it came about is a question we shall
address in Chapter 34.
The Mechanical Layering of the Earth
0
Depth (kilometers)
200
Temperature (°C)
2000
3000
Asthenosphere —
the zone of
partial melting
Molten
mantle
rock
400
600
c
Melting
omplete
begins
800
Solid
mantle
rock
Melting
Mechanical layering means layering in which
mechanical properties, like rigidity (stiffness or elasticity), change abruptly across a boundary even though
chemical composition may not. The only boundary of
that sort we have encountered thus far is the one
between the inner core and the outer core; on the inside
of the boundary matter is solid (rigid) and on the outside
it is liquid (nonrigid), but the composition is essentially
constant across it. As one consequence of that change
in mechanical properties, convection currents are generated in the outer core, with the result that the earth has
a magnetic field. There is yet one other mechanical
boundary with consequences just as profound as the
magnetic field.
About 70 kilometers under the surface (the depth
varies somewhat from place to place), seismic waves
encounter a layer in which they suddenly travel more
slowly than in either the rocks above or those below.
This layer extends to a depth of approximately 250 kilometers and is sometimes called the low velocity zone by
seismologists. The top of the low velocity zone is in the
upper mantle and is not a boundary between different
types of rock (as is, for example, the Mohorovicic discontinuity between the crust and the mantle). Rather, it
is a boundary between solid peridotite and peridotite
1000
Figure 30.9. The white curve shows how temperature
and pressure are related in the earth. The shaded band
is the range of temperatures and pressures within which
peridotite begins to melt. Between depths of about 70
and 250 kilometers, the temperature is within the region
of partial melting of peridotite. This results in the low
velocity zone, or asthenosphere.
297
ing the variation in the velocities of P and S waves as
they traverse these layers. Notice particularly the
decrease in velocity for both types of waves just below
the lithosphere and the absence of S waves beyond the
lower mantle.
100
Lithosphere
Depth (kilometers)
Upper
mantle
Asth
e
nos
ph
ere
Lower
mantle
Isostasy
2900
It is obvious to you that the continents are higher in
elevation than the ocean basins, and part of the reason is
that the oceanic lithosphere is denser and therefore
floats deeper in the underlying asthenosphere.
However, that is not the entire explanation. Seismic
observations show that the continental crust is a few to
several times thicker than the oceanic crust. In fact, the
thickest parts of the continents are where the highest
mountains are.
Recall in Chapter 6 that we discussed buoyancy
and Archimedes’ Principle, and we learned that objects
that float in fluids displace an amount of fluid that
weighs the same as they do. For this reason, icebergs
float only partially immersed in water, and thicker icebergs protrude higher above the water than thinner ones.
This principle is not restricted to objects floating in
water. One can think of the crust as “sinking” until it
has displaced a weight of mantle equal to its own
weight. The oceanic crust is, on the whole, denser and
thinner, so it tends to float lower. The continental rocks
are less dense and thicker, and so float higher. Fold
mountain ranges on the continents are very thick, and so
they are also quite high.
This concept of gravitational equilibrium is called
isostasy. It is simply a special application of the concept of buoyancy. Because the base of the lithosphere
(i.e., top of the asthenosphere) is at a fairly uniform
depth, the concept of isostasy implies that blocks of the
lithosphere that have equal areas on the surface of the
earth also have equal masses.
Figure 30.11 summarizes the relationships among
continental crust, oceanic crust, mantle, lithosphere, and
asthenosphere.
Outer
core
5100
Inner
core
Wave speed (kilometers/second)
14
12
P
s
ve
wa
P waves
s
ave
10
8
Pw
ves
S wa
6
4
2
1000
2000
3000
4000
Depth (kilometers)
5000
6000
Figure 30.10. The top of this figure shows a cross section through the earth, with the various mechanically
distinct layers indicated by different shades of gray.
Below is a graph showing how the velocities of P and S
waves vary with depth in the earth. Notice that the
change from one mechanical layer to another is signaled
by a relatively abrupt change in seismic velocities.
band with increase in depth, peridotite is again solid,
owing to the higher pressure, even though the temperature in the earth is higher.
Floating on this partially molten asthenosphere is
the lithosphere, the rigid outer shell of the earth that we
have identified previously but have never precisely
defined. Now we can recognize that the lithosphere
consists of all the solid earth above the asthenosphere—
the entire thickness of the crust plus the outermost part
of the mantle that is too cool to be partially molten. The
continents are thus part of the lithosphere, and we shall
find later that the lithosphere is cracked into several
large segments that jostle about and rub against one
another.
A cross section of the earth showing the various
mechanically distinct layers is shown at the top of
Figure 30.10. Below the cross section is a graph depict-
Crust
Mantle
Lithosphere
Asthenosphere
Figure 30.11. The relationships among continental
crust, oceanic crust, mantle, lithosphere, and asthenosphere, approximately to the correct vertical scale.
298
Summary
The earth is a differentiated planet, consisting of
layers that become progressively less dense proceeding
from the center out. The densest part, a solid iron (or
iron-nickel) inner core, is surrounded by a liquid outer
core of nearly the same composition. Outside of that is
the mantle, which consists mostly of oxygen, silicon,
magnesium, and iron, these atoms forming compounds
called dense oxides in the lower mantle and forming the
rock called peridotite in the upper mantle. Surrounding
the mantle is a thin crust consisting of granitic rock
(continental crust) or basalt (oceanic crust). The division into core-mantle-crust is a chemical differentiation,
and the division of the core as solid and liquid is
mechanical (i.e., the mechanical properties change
across the boundary, but the chemical composition does
not). Another mechanical division differentiates the
lithosphere, a brittle outer shell consisting of the crust
and cool uppermost mantle, from the asthenosphere, a
zone below the lithosphere in which the peridotite of the
mantle is partially molten and plastic. The existence of
the lithosphere and the asthenosphere has a profound
effect on the way the earth works, and we shall now
begin to consider that.
6.
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13.
STUDY GUIDE
Chapter 30: The Interior of the Earth
14.
A. FUNDAMENTAL PRINCIPLES: No new fundamental principles.
15.
B. MODELS, IDEAS, QUESTIONS, OR APPLICATIONS
1. Is it possible to “weigh” the earth?
2. What can be deduced by studying the “magnetic
effects” of the earth?
3. What can be learned about the interior of the earth
by “listening” to earthquakes?
4. What is the current view of the structure, composition, density, and other physical properties of the
earth’s interior?
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C. GLOSSARY
1. Asthenosphere: A zone in the upper mantle where
the peridotite is partially molten so it can be
deformed and flow. (The asthenosphere is also
referred to as the low velocity zone.)
2. Atmosphere: See Chapter 28.
3. Cavendish Balance: Equipment used to determine
relative densities and the universal gravitational
constant, G.
4. Crust: The uppermost layer of the earth, consisting of silicates of two general types: continental
and oceanic crust.
5. Curie Temperature: A particular elevated tem-
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perature above which a magnetic substance loses
its bulk magnetic properties. The Curie temperature of iron is 760 °C.
Dense Oxide: Compounds which have the same
chemical composition as peridotite but, due to the
temperatures and pressures of the lower mantle,
have their atoms rearranged by chemical reactions
to form denser, more stable compounds.
Fault: A fracture in a rock structure along which
portions of the earth have moved relative to one
another.
Hydrosphere: See Chapter 28.
Hypocenter: The point in the subsurface of the
earth where the disturbance (earthquake) actually
occurred. The point on the surface of the earth
above the hypocenter is called the epicenter.
Inner Core: The layer of the earth directly underneath the liquid outer core consisting of solid iron
and nickel.
Isostasy: Gravitational equilibrium (balance of
forces) of the earth’s crust. The crust sinks until it
displaces a weight of mantle equal to its own
weight and thus “floats” in the mantle at rest.
Latent Heat of Crystallization: Heat released
when liquid matter crystallizes into its solid form.
Lithosphere: The rigid outer shell of the earth
which consists of the crust and the outermost part
of the mantle that is too cool to be partially molten.
Mantle: The layer of the earth directly underneath
the crust, consisting of peridotite in the upper mantle and dense oxides in the lower mantle.
Mechanical Layering:
Layering in which
mechanical properties, such as rigidity, change
abruptly across a boundary even though chemical
composition may not.
Outer Core: The layer of the earth directly underneath the mantle, consisting of molten (liquid) iron
and nickel.
Peridotite: A rock type that makes up the upper
mantle; it contains less silicon and more magnesium than is found in the rocks of the crust.
Primary Wave: Often shortened to P waves, they
are compression seismic waves and arrive at the
remote detectors (seismometers) first.
Secondary Wave: Often shortened to S waves,
they are shear seismic waves and arrive at the
remote detectors (seismometers) after the primary
waves.
Seismic Discontinuity: Any depth in the earth at
which seismic wave velocities experience a distinct
and abrupt change. The Mohorovicic discontinuity (Moho) is a seismic discontinuity that marks the
boundary between the crust and the mantle.
Seismic Wave: Shock waves generated in the
earth when rock breaks abruptly during an earthquake.
22. Seismometer: An earthquake-measuring device.
The graphical record of the earthquake produced by
a seismometer is called a seismograph.
23. Shadow Zone: A zone in which there is little or no
seismic wave detection from a given earthquake
event. It extends in a 40°-wide band around the
earth, 103° to 143° from the epicenter of the earthquake; the locations of the shadow zones are different for all earthquakes, if the locations of the epicenters are different.
24. Silicate Rock: A rock made of minerals which
contain the SiO44– molecular ion. Granite and
basalt are examples of silicate rocks.
30.7. From center to surface, the chemically distinct layers of the earth are
(a) inner core, outer core, mantle, crust.
(b) core, mantle, lithosphere.
(c) inner core, outer core, lower mantle, upper
mantle, crust.
(d) core, mantle, crust.
30.8. Why are fold mountain belts topographically
high?
30.9. The asthenosphere consists of
(a) completely molten peridotite.
(b) partially molten peridotite.
(c) completely solid peridotite.
(d) rock that is either granitic or basalt, depending
on location.
D. FOCUS QUESTIONS
1. Consider what has been learned from seismic
waves about the interior of the earth:
a. Sketch curves showing the changes in speed of
both primary and secondary waves as a function of
depth below the earth’s surface.
b. Sketch the interior layers of the earth showing
the relationship of these layers to the observed
wave speeds. Describe the physical characteristics
of each layer.
c. Describe the shadow zones for both primary
and secondary waves and explain why they occur.
E. EXERCISES
30.1. From the average density of the earth, we
know that
(a) the interior of the earth is of uniform density.
(b) density increases with depth in the earth.
(c) the earth is a layered planet.
(d) density decreases with depth in the earth.
30.2. Seismic P waves are
(a) compressional waves.
(b) shear waves.
(c) surface waves.
(d) none of the above.
30.3. Seismic waves travel in curved paths through
the earth but make abrupt changes in direction and
speed at seismic discontinuities. Why?
30.4. What are the principal seismic discontinuities
in the earth?
30.5. If the outer core and inner core are of essentially the same composition, why is the inner core solid?
30.6. The earth is a differentiated planet. What is
meant by that?
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