Survey
* Your assessment is very important for improving the workof artificial intelligence, which forms the content of this project
* Your assessment is very important for improving the workof artificial intelligence, which forms the content of this project
Schiehallion experiment wikipedia , lookup
History of Earth wikipedia , lookup
Late Heavy Bombardment wikipedia , lookup
History of geology wikipedia , lookup
Plate tectonics wikipedia , lookup
Age of the Earth wikipedia , lookup
Mantle plume wikipedia , lookup
Large igneous province wikipedia , lookup
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 291 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 292 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 293 (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 294 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 295 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. 7. 8. 9. 10. 11. 12. 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? 16. 17. 18. 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- 19. 20. 21. 299 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? 300