6.4: Earth's Interior

Understanding Earth's Interior Using Seismology

Before you can learn about plate tectonics, you need to know something about the layers that are found inside Earth. These layers are divided by composition into core, mantle, and crust or by mechanical properties into lithosphere and asthenosphere. Scientists use information from earthquakes and computer modeling to learn about Earth’s interior.

EXPLORING EARTH’S INTERIOR

How do scientists know what is inside the Earth? We don’t have direct evidence! Rocks yield some clues, but they only reveal information about the outer crust. In rare instances, a mineral, such as a diamond, comes to the surface from deeper down in the crust or the mantle. To learn about Earth’s interior, scientists use energy to “see” the different layers of the Earth, just like doctors can use an MRI, CT scan, or x-ray to see inside our bodies.

Seismic Waves

P-waves and S-waves move through different parts of the earth’s interior in different ways. Analogous to how you can see what is in the room around you by interpreting the light that your eyes receive, light that has interacted with the things in the room around you to give it its characteristics, seismologists can interpret recordings of seismic waves to “see” inside the earth. Such imaging of the earth’s interior is based on how the different layers of the earth have affected the seismic waves in different ways.

Where seismic waves speed up or slow down, they refract, changing the direction in which they are traveling. Where seismic waves encounter an abrupt boundary between two very different layers, some of the seismic wave energy is reflected, bouncing back at the same angle it struck. The reflections and refractions of seismic waves allow the layers and boundaries within the earth to be located and studied.

Here are some examples of what we have been able to distinguish in the earth’s interior from the study of seismic waves and how they travel through the layers of the earth:

The thickness of the crust. This is a measure of the thickness of the crust based on the abrupt increase in speed of seismic waves that occurs when they enter the mantle. The boundary between the crust and mantle, as inferred from the change in the speed of P- and S-waves, is called the Mohorovicic discontinuity, named after the Croatian seismologist who first discerned it; usually it is referred to simply as the Moho. It is mainly from seismic waves that we know how thin oceanic crust is and how thick continental crust is.
The thickness of the lithosphere. Where seismic waves pass down from the lithosphere into the asthenosphere, they slow down. This is because of the lower rigidity and compressibility of the rocks in the layer below the lithosphere. The zone below the lithosphere where seismic waves travel more slowly is called the low velocity zone. The low velocity zone is probably coincident with the asthenosphere.
The boundary between the upper and lower mesosphere (upper and lower mantle). This shows up as an increase in seismic wave speed at a depth of 660 km.
The boundary between the mantle and the core. This is marked by S-waves coming to an abrupt stop, presumably because the outer core is liquid, and a sudden large reduction in the speed of P-waves, as they enter the liquid core where there is no rigidity to contribute to P-wave speed.
The inner core. This was first recognized by refraction of P-waves passing through this part of the core, due to an abrupt increase in their speed, which was not shown by P-waves traveling through only the outer part of the core.
Seismic tomography: imaging slabs and masses at various orientations in the earth, not just in layers. By combining data from many seismometers, three-dimensional images of zones in the earth that have higher or lower seismic wave speeds can be constructed. Seismic tomography shows that in some places there are masses of what may be subducted plates that have penetrated below the asthenosphere into the mesosphere and, in some cases, penetrated into the lower mesosphere, the deepest part of the mantle. In other places, subducted plates appear to have piled up at the base of the upper mesosphere without penetrating into the lower mesosphere.

One ingenious way scientists learn about Earth’s interior is by looking at how energy travels from the point of an earthquake. These are seismic waves (figure 1). Seismic waves travel outward in all directions from where the ground breaks at an earthquake. These waves are picked up by seismographs around the world. Two types of seismic waves are most useful for learning about Earth’s interior.

P-waves (primary waves) are fastest, traveling at about 6 to 7 kilometers (about 4 miles) per second, so they arrive first at the seismometer. P-waves move in a compression/expansion type motion, squeezing and unsqueezing earth materials as they travel. This produces a change in volume for the material. P-waves bend slightly when they travel from one layer into another. Seismic waves move faster through denser or more rigid material. As P-waves encounter the liquid outer core, which is less rigid than the mantle, they slow down. This makes the P-waves arrive later and further away than would be expected. The result is a P-wave shadow zone. No P-waves are picked up at seismographs 104° to 140° from the earthquakes focus.

Figure 1. How P-waves travel through Earth’s interior.
S-waves (secondary waves) are about half as fast as P-waves, traveling at about 3.5 km (2 miles) per second, and arrive second at seismographs. S-waves move in an up and down motion perpendicular to the direction of wave travel. This produces a change in shape for the earth materials they move through. Only solids resist a change in shape, so S-waves are only able to propagate through solids. S-waves cannot travel through liquid.

By tracking seismic waves, scientists have learned what makes up the planet’s interior (figure 2).

P-waves slow down at the mantle core boundary, so we know the outer core is less rigid than the mantle.
S-waves disappear at the mantle core boundary, so the outer core is liquid.

Figure 2. Letters describe the path of an individual P-wave or S-wave. Waves traveling through the core take on the letter K.

This animation shows a seismic wave shadow zone.

Other Clues About Earth's Interior

Earth’s overall density is higher than the density of crustal rocks, so the core must be made of something dense, like metal.
Since Earth has a magnetic field, there must be metal within the planet. Iron and nickel are both magnetic.
Meteorites are the remains of the material that formed the early solar system and are thought to be similar to material in Earth’s interior (figure 3).

Figure 3. This meteorite contains silica minerals and iron-nickel. The material is like the boundary between Earth’s core and mantle.

INTRODUCTION

Three centuries ago, the English scientist Isaac Newton calculated, from his studies of planets and the force of gravity, that the average density of the Earth is twice that of surface rocks and therefore that the Earth’s interior must be composed of much denser material. Our knowledge of what’s inside the Earth has improved immensely since Newton’s time, but his estimate of the density remains essentially unchanged. Our current information comes from studies of the paths and characteristics of earthquake waves travelling through the Earth, as well as from laboratory experiments on surface minerals and rocks at high pressure and temperature. Other important data on the Earth’s interior come from geological observation of surface rocks and studies of the Earth’s motions in the Solar System, its gravity and magnetic fields, and the flow of heat from inside the Earth.The planet Earth is made up of three main shells: the very thin, brittle crust, the mantle, and the core; the mantle and core are each divided into two parts. All parts are drawn to scale on the cover of this publication, and a table at the end lists the thicknesses of the parts. Although the core and mantle are about equal in thickness, the core actually forms only 15 percent of the Earth’s volume, whereas the mantle occupies 84 percent. The crust makes up the remaining 1 percent. Our knowledge of the layering and chemical composition of the Earth is steadily being improved by earth scientists doing laboratory experiments on rocks at high pressure and analyzing earthquake records on computers.

Earth’s Interior

The interior of a planet—even our own Earth—is difficult to study, and its composition and structure must be determined indirectly. Our only direct experience is with the outermost skin of Earth^’s crust, a layer no more than a few kilometers deep. It is important to remember that, in many ways, we know less about our own planet 5 kilometers beneath our feet than we do about the surfaces of Venus and Mars.

Earth is composed largely of metal and silicate rock (see the Composition and Structure of Planets section). Most of this material is in a solid state, but some of it is hot enough to be molten. The structure of material in Earth’s interior has been probed in considerable detail by measuring the transmission of seismic waves through Earth. These are waves that spread through the interior of Earth from earthquakes or explosion sites.

Seismic waves travel through a planet rather like sound waves through a struck bell. Just as the sound frequencies vary depending on the material the bell is made of and how it is constructed, so a planet^’s response depends on its composition and structure. By monitoring the seismic waves in different locations, scientists can learn about the layers through which the waves have traveled. Some of these vibrations travel along the surface; others pass directly through the interior. Seismic studies have shown that Earth’s interior consists of several distinct layers with different compositions, illustrated in Figure 8.3. As waves travel through different materials in Earth’s interior, the waves—just like light waves in telescope lenses—bend (or refract) so that some seismic stations on Earth receive the waves and others are in “shadows.” Detecting the waves in a network of seismographs helps scientists construct a model of Earth’s interior, showing liquid and solid layers. This type of seismic imaging is not unlike that used in ultrasound, a type of imaging used to see inside the body.

Cut-away View of the Interior of the Earth. This illustration shows the globe of the Earth with a wedge-shaped portion removed to reveal the interior. The inner core is labeled and represented as a small yellow sphere at the center. Next, the core is shown in orange and surrounds the inner core. The larger mantle surrounds the core and is drawn in taupe. Finally, the crust is indicated as a thin blue line. — Figure 8.3 : Interior Structure of Earth. The crust, mantle, and inner and outer cores (solid and liquid, respectively) as shown as revealed by seismic studies.

The top layer is the crust, the part of Earth we know best (Figure 8.4). Oceanic crust covers 55% of Earth’s surface and lies mostly submerged under the oceans. It is typically about 6 kilometers thick and is composed of volcanic rocks called basalt. Produced by the cooling of volcanic lava, basalts are made primarily of the elements silicon, oxygen, iron, aluminum, and magnesium. The continental crust covers 45% of the surface, some of which is also beneath the oceans. The continental crust is 20 to 70 kilometers thick and is composed predominantly of a different volcanic class of silicates (rocks made of silicon and oxygen) called granite. These crustal rocks, both oceanic and continental, typically have densities of about 3 g/cm³. (For comparison, the density of water is 1 g/cm³.) The crust is the easiest layer for geologists to study, but it makes up only about 0.3% of the total mass of Earth.

Computer-generated image of the entire Earth’s crust, including the details of the ocean floor. — Figure 8.4 : Earth’s Crust. This computer-generated image shows the surface of Earth’s crust as determined from satellite images and ocean floor radar mapping. Oceans and lakes are shown in blue, with darker areas representing depth. Dry land is shown in shades of green and brown, and the Greenland and Antarctic ice sheets are depicted in shades of white. (credit: modification of work by C. Amante, B. W. Eakins, National Geophysical Data Center, NOAA)

The largest part of the solid Earth, called the mantle, stretches from the base of the crust downward to a depth of 2900 kilometers. The mantle is more or less solid, but at the temperatures and pressures found there, mantle rock can deform and flow slowly. The density in the mantle increases downward from about 3.5 g/cm³ to more than 5 g/cm³ as a result of the compression produced by the weight of overlying material. Samples of upper mantle material are occasionally ejected from volcanoes, permitting a detailed analysis of its chemistry.

Beginning at a depth of 2900 kilometers, we encounter the dense metallic core of Earth. With a diameter of 7000 kilometers, our core is substantially larger than the entire planet Mercury. The outer core is liquid, but the innermost part of the core (about 2400 kilometers in diameter) is probably solid. In addition to iron, the core probably also contains substantial quantities of nickel and sulfur, all compressed to a very high density.

The separation of Earth into layers of different densities is an example of differentiation, the process of sorting the major components of a planet by density. The fact that Earth is differentiated suggests that it was once warm enough for its interior to melt, permitting the heavier metals to sink to the center and form the dense core. Evidence for differentiation comes from comparing the planet’s bulk density (5.5 g/cm³) with the surface materials (3 g/cm³) to suggest that denser material must be buried in the core.

THE CRUST

Let us now examine our planet’s outer layers in more detail. Earth’s crust is a dynamic place. Volcanic eruptions, erosion, and large-scale movements of the continents rework the surface of our planet constantly. Geologically, ours is the most active planet. Many of the geological processes described in this section have taken place on other planets as well, but usually in their distant pasts. Some of the moons of the giant planets also have impressive activity levels. For example, Jupiter’s moon Io has a remarkable number of active volcanoes.

Composition of the Crust

Earth’s crust is largely made up of oceanic basalt and continental granite. These are both igneous rock, the term used for any rock that has cooled from a molten state. All volcanically produced rock is igneous (Figure 8.6).

Image of a Lava Flow from a Basaltic Eruption. The leading edge of the flow is red-hot, while the surface behind has cooled to almost black — Figure 8.6 : Formation of Igneous Rock as Liquid Lava Cools and Freezes. This is a lava flow from a basaltic eruption. Basaltic lava flows quickly and can move easily over distances of more than 20 kilometers. (credit: USGS)

Two other kinds of rock are familiar to us on Earth, although it turns out that neither is common on other planets. Sedimentary rocks are made of fragments of igneous rock or the shells of living organisms deposited by wind or water and cemented together without melting. On Earth, these rocks include the common sandstones, shales, and limestones. Metamorphic rocks are produced when high temperature or pressure alters igneous or sedimentary rock physically or chemically (the word metamorphic means “changed in form”). Metamorphic rocks are produced on Earth because geological activity carries surface rocks down to considerable depths and then brings them back up to the surface. Without such activity, these changed rocks would not exist at the surface.

There is a fourth very important category of rock that can tell us much about the early history of the planetary system: primitive rock, which has largely escaped chemical modification by heating. Primitive rock represents the original material out of which the planetary system was made. No primitive material is left on Earth because the entire planet was heated early in its history. To find primitive rock, we must look to smaller objects such as comets, asteroids, and small planetary moons. We can sometimes see primitive rock in samples that fall to Earth from these smaller objects.

A block of quartzite on Earth is composed of materials that have gone through all four of these states. Beginning as primitive material before Earth was born, it was heated in the early Earth to form igneous rock, transformed chemically and redeposited (perhaps many times) to form sedimentary rock, and finally changed several kilometers below Earth’s surface into the hard, white metamorphic stone we see today.

Because the crust is accessible to us, its geology has been extensively studied, and therefore much more information is known about its structure and composition than about the structure and composition of the mantle and core. Within the crust, intricate patterns are created when rocks are redistributed and deposited in layers through the geologic processes of eruption and intrusion of lava, erosion, and consolidation of rock particles, and solidification and recrystallization of porous rock.

Figure 1. The oceanic crust at the island of Hawaii is about 5 kilometers thick. The thickness of the continental crust under eastern California ranges from 25 kilometers under the Great Valley to 60 kilometers under the Sierra Nevada.

By the large-scale process of plate tectonics, about twelve plates, which contain combinations of continents and ocean basins, have moved around on the Earth’s surface through much of geologic time. The edges of the plates are marked by concentrations of earthquakes and volcanoes. Collisions of plates can produce mountains like the Himalayas, the tallest range in the world. The plates include the crust and part of the upper mantle, and they move over a hot, yielding upper mantle zone at very slow rates of a few centimeters per year, slower than the rate at which fingernails grow. The crust is much thinner under the oceans than under continents (see figure above).

The boundary between the crust and mantle is called the Mohorovicic discontinuity (or Moho); it is named in honor of the man who discovered it, the Croatian scientist Andrija Mohorovicic. No one has ever seen this boundary, but it can be detected by a sharp increase downward in the speed of earthquake waves there. The explanation for the increase at the Moho is presumed to be a change in rock types. Drill holes to penetrate the Moho have been proposed, and a Soviet hole on the Kola Peninsula has been drilled to a depth of 12 kilometers, but drilling expense increases enormously with depth, and Moho penetration is not likely very soon.

THE MANTLE

Our knowledge of the upper mantle, including the tectonic plates, is derived from analyses of earthquake waves (see figure for paths); heat flow, magnetic, and gravity studies; and laboratory experiments on rocks and minerals. Between 100 and 200 kilometers below the Earth’s surface, the temperature of the rock is near the melting point; molten rock erupted by some volcanoes originates in this region of the mantle. This zone of extremely yielding rock has a slightly lower velocity of earthquake waves and is presumed to be the layer on which the tectonic plates ride. Below this low-velocity zone is a transition zone in the upper mantle; it contains two discontinuities caused by changes from less dense to more dense minerals. The chemical composition and crystal forms of these minerals have been identified by laboratory experiments at high pressure and temperature. The lower mantle, below the transition zone, is made up of relatively simple iron and magnesium silicate minerals, which change gradually with depth to very dense forms. Going from mantle to core, there is a marked decrease (about 30 percent) in earthquake wave velocity and a marked increase (about 30 percent) in density.

Figure 2. Cross section of the whole Earth, showing the complexity of paths of earthquake waves. The paths curve because the different rock types found at different depths change the speed at which the waves travel. Solid lines marked P are compressional waves; dashed lines marked S are shear waves. S waves do not travel through the core but may be converted to compressional waves (marked K) on entering the core (PKP, SKS). Waves may be reflected at the surface (PP, PPP, SS).

The core was the first internal structural element to be identified. It was discovered in 1906 by R.D. Oldham, from his study of earthquake records, and it helped to explain Newton’s calculation of the Earth’s density. The outer core is presumed to be liquid because it does not transmit shear (S) waves and because the velocity of compressional (P) waves that pass through it is sharply reduced. The inner core is considered to be solid because of the behavior of P and S waves passing through it.

Data from earthquake waves, rotations and inertia of the whole Earth, magnetic-field dynamo theory, and laboratory experiments on melting and alloying of iron all contribute to the identification of the composition of the inner and outer core. The core is presumed to be composed principally of iron, with about 10 percent alloy of oxygen or sulfur or nickel, or perhaps some combination of these three elements.

This table of depths, densities, and composition is derived mostly from information in a textbook by Don L. Anderson. Scientists are continuing to refine the chemical and mineral composition of the Earth’s interior by laboratory experiments, by using pressures 2 million times the pressure of the atmosphere at the surface and temperatures as high as 20000C.

Table 1. Data on the Earth’s Interior
	Thickness (km)	Density (g/cm)		Types of Rock Found
	Thickness (km)	Top	Bottom	Types of Rock Found
Crust	30	2.2	2.9	Silicic rocks. Andesite, basalt at base.
Upper mantle	720	3.4	4.4	Peridotite, eclogite, olivine, spinel, garnet, pyroxene. Perovskite, oxides
Lower mantle	2,171	4.4	5.6	Magnesium and silicon oxides.
Outer core	2,259	9.9	12.1	Iron + oxygen, sulfur, nickel alloy.
Inner Core	1,221	12.8	13.1	Iron + oxygen, sulfur, nickel alloy.
Total Thickness	6,401

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THE EARTH’S LAYERS

The layers scientists recognize are pictured in figure 4.

Figure 4. A cross section of Earth showing the following layers: (1) crust (2) mantle (3a) outer core (3b) inner core (4) lithosphere (5) asthenosphere (6) outer core (7) inner core.

Core, mantle, and crust are divisions based on composition:

The crust is less than 1% of Earth by mass. The oceanic crust is mafic, whilecontinental crust is often more felsic rock.
The mantle is hot, ultramafic rock. It represents about 68% of Earth’s mass.
The core is mostly iron metal. The core makes up about 31% of the Earth.

Lithosphere and asthenosphere are divisions based on mechanical properties:

The lithosphere is composed of both the crust and the portion of the upper mantle that behaves as a brittle, rigid solid.
The asthenosphere is partially molten upper mantle material that behaves plastically and can flow.

This animation shows the layers by composition and by mechanical properties.

Crust and Lithosphere

Earth’s outer surface is its crust; a cold, thin, brittle outer shell made of rock. The crust is very thin, relative to the radius of the planet. There are two very different types of crust, each with its own distinctive physical and chemical properties, which are summarized in table 1.

Table 1.
Crust	Thickness	Density	Composition	Rock types
Oceanic	5-12 km (3-8 mi)	3.0 g/cm³	Mafic	Basalt and gabbro
Continental	Avg. 35 km (22 mi)	2.7 g/cm³	Felsic	All types

Oceanic crust is composed of mafic magma that erupts on the seafloor to create basalt lava flows or cools deeper down to create the intrusive igneous rock gabbro (figure 5).

Sediments, primarily muds and the shells of tiny sea creatures, coat the seafloor. Sediment is thickest near the shore where it comes off the continents in rivers and on wind currents.

Continental crust is made up of many different types of igneous, metamorphic, and sedimentary rocks. The average composition is granite, which is much less dense than the mafic rocks of the oceanic crust (figure 6). Because it is thick and has relatively low density, continental crust rises higher on the mantle than oceanic crust, which sinks into the mantle to form basins. When filled with water, these basins form the planet’s oceans.

Figure 6. This granite from Missouri is more than 1 billion years old.

The lithosphere is the outermost mechanical layer, which behaves as a brittle, rigid solid. The lithosphere is about 100 kilometers thick. Look at figure 4. Can you find where the crust and the lithosphere are located? How are they different from each other?

The definition of the lithosphere is based on how earth materials behave, so it includes the crust and the uppermost mantle, which are both brittle. Since it is rigid and brittle, when stresses act on the lithosphere, it breaks. This is what we experience as an earthquake.

Mantle

The two most important things about the mantle are: (1) it is made of solid rock, and (2) it is hot. Scientists know that the mantle is made of rock based on evidence from seismic waves, heat flow, and meteorites.

The properties fit the ultramafic rock peridotite, which is made of the iron- and magnesium-rich silicate minerals (figure 7). Peridotite is rarely found at Earth’s surface.

Figure 7. Peridotite is formed of crystals of olivine (green) and pyroxene (black).

Scientists know that the mantle is extremely hot because of the heat flowing outward from it and because of its physical properties.

Heat flows in two different ways within the Earth:

Conduction: Heat is transferred through rapid collisions of atoms, which can only happen if the material is solid. Heat flows from warmer to cooler places until all are the same temperature. The mantle is hot mostly because of heat conducted from the core.
Convection: If a material is able to move, even if it moves very slowly, convection currents can form.

Convection in the mantle is the same as convection in a pot of water on a stove. Convection currents within Earth’s mantle form as material near the core heats up. As the core heats the bottom layer of mantle material, particles move more rapidly, decreasing its density and causing it to rise. The rising material begins the convection current. When the warm material reaches the surface, it spreads horizontally. The material cools because it is no longer near the core. It eventually becomes cool and dense enough to sink back down into the mantle. At the bottom of the mantle, the material travels horizontally and is heated by the core. It reaches the location where warm mantle material rises, and the mantle convection cell is complete (figure 8).

Core

At the planet’s center lies a dense metallic core. Scientists know that the core is metal because:

The density of Earth’s surface layers is much less than the overall density of the planet, as calculated from the planet’s rotation. If the surface layers are less dense than average, then the interior must be denser than average. Calculations indicate that the core is about 85% iron metal with nickel metal making up much of the remaining 15%.
Metallic meteorites are thought to be representative of the core. The 85% iron/15% nickel calculation above is also seen in metallic meteorites (figure 9).

Figure 9. An iron meteorite is the closest thing to the Earth’s core that we can hold in our hands.

If Earth’s core were not metal, the planet would not have a magnetic field. Metals such as iron are magnetic, but rock, which makes up the mantle and crust, is not.

Scientists know that the outer core is liquid and the inner core is solid because:

S-waves stop at the inner core.
The strong magnetic field is caused by convection in the liquid outer core. Convection currents in the outer core are due to heat from the even hotter inner core.

The heat that keeps the outer core from solidifying is produced by the breakdown of radioactive elements in the inner core.

LESSON SUMMARY

Earth is made of three layers: the crust, mantle, and core.
The brittle crust and uppermost mantle are together called the lithosphere.
Beneath the lithosphere, the mantle is solid rock that can flow, or behave plastically.
The hot core warms the base of the mantle, which causes mantle convection.

Seismology is the study of vibrations within Earth. These vibrations are caused by various events: earthquakes, extraterrestrial impacts, explosions, storm waves hitting the shore, and tidal effects. Of course, seismic techniques have been most widely applied to the detection and study of earthquakes, but there are many other applications, and arguably seismic waves provide the most important information that we have concerning Earth’s interior. Before going any deeper into Earth, however, we need to take a look at the properties of seismic waves. The types of waves that are useful for understanding Earth’s interior are called body waves, meaning that, unlike the surface waves on the ocean, they are transmitted through Earth materials.

Figure \(\PageIndex{1}\) Hitting a large block of rock with a heavy hammer will create seismic waves within the rock. Please don’t try this at home! © Steven Earle. CC BY.

Imagine hitting a large block of strong rock (e.g., granite) with a heavy sledgehammer (Figure \(\PageIndex{1}\)). At the point where the hammer strikes it, a small part of the rock will be compressed by a fraction of a millimetre. That compression will transfer to the neighbouring part of the rock, and so on through to the far side of the rock—all in a fraction of a second. This is known as a compression wave, and it can be illustrated by holding a loose spring (like a Slinky) that is attached to something (or someone) at the other end. If you give it a sharp push so the coils are compressed, the compression propagates (travels) along the length of the spring and back (Figure \(\PageIndex{2}\)). You can think of a compression wave as a “push” wave—it’s called a P wave (although the “P” stands for “primary” because P waves arrive first at seismic stations).

When we hit a rock with a hammer, we also create a different type of body wave, one that is characterized by back-and-forth vibrations (as opposed to compressions). This is known as a shear wave (S wave, where the “S” stands for “secondary”), and an analogy would be what happens when you flick a length of rope with an up-and-down motion. As shown in Figure \(\PageIndex{2}\), a wave will form in the rope, which will travel to the end of the rope and back.

Figure \(\PageIndex{2}\) A compression wave can be illustrated by a spring (like a Slinky) that is given a sharp push at one end. A shear wave can be illustrated by a rope that is given a quick flick. © Steven Earle. CC BY.

Compression waves and shear waves travel very quickly through geological materials. As shown in Figure \(\PageIndex{3}\), typical P wave velocities are between 0.5 kilometers per second (km/s) and 2.5 km/s in unconsolidated sediments, and between 3.0 km/s and 6.5 km/s in solid crustal rocks. Of the common rocks of the crust, velocities are greatest in basalt and granite. S waves are slower than P waves, with velocities between 0.1 km/s and 0.8 km/s in soft sediments, and between 1.5 km/s and 3.8 km/s in solid rocks.

Figure \(\PageIndex{3}\) Typical velocities of P-waves (red) and S-waves (blue) in sediments and in solid crustal rocks. “P Wave Velocity, m/s” and “Shear Wave Velocity, m/s” by the US Environment Protection Agency. Edited by Steven Earle. Public domain.

Exercise 9.1 How soon will seismic waves get here?

Imagine that a strong earthquake takes place on Vancouver Island within Strathcona Park (west of Courtenay). Assuming that the crustal average P wave velocity is 5 km per second, how long will it take (in seconds) for the first seismic waves (P waves) to reach you in the following places (distances from the epicentre are shown)?

Nanaimo (120 km away)
Surrey (200 km away)
Kamloops (390 km away)

See Appendix 3 for Exercise 9.1 answers.

Mantle rock is generally denser and stronger than crustal rock and both P- and S-waves travel faster through the mantle than they do through the crust. Moreover, seismic-wave velocities are related to how tightly compressed a rock is, and the level of compression increases dramatically with depth. Finally, seismic waves are affected by the phase state of rock. They are slowed if there is any degree of melting in the rock. If the material is completely liquid, P waves are slowed dramatically and S waves are stopped altogether.

Figure \(\PageIndex{4}\) P wave (red) and S wave (blue) velocity variations with depth in Earth. The diagram on the right shows an expanded view of the upper 660 kilometers of the curves in the diagram on the left. © Steven Earle. CC BY.

As shown on the right-hand part of Figure \(\PageIndex{4}\), the upper approximately 100 km of the Earth is known as the lithosphere. This includes the rigid upper part of the mantle (or lithospheric mantle) and the crust. The next 150 km is the asthenosphere or low velocity zone (because seismic waves are slowed as they pass through that material). As we’ll see below, that part of the mantle is close to it’s melting point and in some regions may be partially molten.

Accurate seismometers have been used for earthquake studies since the late 1800s, and systematic use of seismic data to understand Earth’s interior started in the early 1900s. The rate of change of seismic waves with depth in Earth (as shown in Figure \(\PageIndex{4}\)) has been determined over the past several decades by analyzing seismic signals from large earthquakes at seismic stations around the world. Small differences in arrival time of signals at different locations have been interpreted to show that:

Velocities are greater in mantle rock than in the crust.
Velocities generally increase with pressure, and therefore with depth.
Velocities slow in the area between a 100 and 250 kilometre depth (called the “low-velocity zone”; equivalent to the asthenosphere).
Velocities increase dramatically at 660 kilometre depth (because of a mineralogical transition).
Velocities slow in the region just above the core-mantle boundary (the D” (d-double-prime) layer or “ultra-low-velocity zone”).
S waves do not pass through the outer part of the core.
P wave velocities increase dramatically at the boundary between the liquid outer core and the solid inner core.

One of the first discoveries about Earth’s interior made through seismology was in 1909 when Croatian seismologist Andrija Mohorovičić (pronounced Moho-ro-vi-chich) realized that at certain distances from an earthquake, two separate sets of seismic waves arrived at a seismic station within a few seconds of each other. He reasoned that the waves that went down into the mantle, traveled through the mantle, and then were bent upward back into the crust, reached the seismic station first because although they had farther to go, they traveled faster through mantle rock (as shown in Figure \(\PageIndex{5}\)). The boundary between the crust and the mantle is known as the Mohorovičić discontinuity (or Moho). Its depth is between 30 and 40 kilometers beneath most of the continental crust, and between 5 and 10 kilometers beneath the oceanic crust.

Figure \(\PageIndex{5}\) Depiction of seismic waves emanating from an earthquake (red star). Some waves travel through the crust to the seismic station (at about 6 km/s), while others go down into the mantle (where they travel at around 8 km/s) and are bent upward toward the surface, reaching the station before the ones that traveled only through the crust. © Steven Earle. CC BY.

Our current understanding of the patterns of seismic wave transmission through Earth is summarized in Figure \(\PageIndex{6}\). Because of the gradual increase in density (and therefore rock strength) with depth, all waves are refracted (toward the lower density material) as they travel through homogenous parts of Earth and thus tend to curve outward toward the surface. Waves are also refracted at boundaries within Earth, such as at the Moho, at the core-mantle boundary (CMB), and at the outer-core/inner-core boundary.

S waves do not travel through liquids—they are stopped at the CMB—and there is an S wave shadow on the side of Earth opposite a seismic source. The angular distance from the seismic source to the shadow zone is 103° on either side, so the total angular distance of the shadow zone is 154°. We can use this information to infer the depth to the CMB.

P waves do travel through liquids, so they can make it through the liquid part of the core. Because of the refraction that takes place at the CMB, waves that travel through the core are bent away from the surface, and this creates a P wave shadow zone on either side, from 103° to 150°. This information can be used to discover the differences between the inner and outer parts of the core.

Figure \(\PageIndex{6}\) Patterns of seismic wave propagation through Earth’s mantle and core. S waves do not travel through the liquid outer core, so they leave a shadow on Earth’s far side where they cannot get to. P waves do travel through the core, but because the waves that enter the core are refracted, there are also P wave shadow zones. © Steven Earle. CC BY.

Exercise 9.2 Liquid Cores in Other Planets

Figure \(\PageIndex{7}\) © Steven Earle. CC BY.

We know that other planets must have (or at least did have) liquid cores like ours, and we could use seismic data to find out how big they are. The S wave shadow zones on planets A and B are shown. Using the same method used for Earth (on the left), sketch in the outlines of the cores for these two other planets.

Using data from many seismometers and hundreds of earthquakes, it is possible to create a two- or three-dimensional image of the seismic properties of part of the mantle. This technique is known as seismic tomography, and an example of the result is shown in Figure \(\PageIndex{8}\).

The Pacific Plate subducts beneath Tonga and appears in Figure \(\PageIndex{8}\) as a 100 kilometre thick slab of cold (blue-colored) oceanic crust that has pushed down into the surrounding hot mantle. The cold rock is more rigid than the surrounding hot mantle rock, so it is characterized by slightly faster seismic velocities. There is volcanism in the Lau spreading centre and also in the Fiji area, and the warm rock in these areas has slower seismic velocities (yellow and red colors).

LESSON SUMMARY

Earth is made of three layers: the crust, mantle, and core.
The brittle crust and uppermost mantle are together called the lithosphere.
Beneath the lithosphere, the mantle is solid rock that can flow, or behave plastically.
The hot core warms the base of the mantle, which causes mantle convection.

Seismology is the study of vibrations within Earth. These vibrations are caused by various events: earthquakes, extraterrestrial impacts, explosions, storm waves hitting the shore, and tidal effects. Of course, seismic techniques have been most widely applied to the detection and study of earthquakes, but there are many other applications, and arguably seismic waves provide the most important information that we have concerning Earth’s interior. Before going any deeper into Earth, however, we need to take a look at the properties of seismic waves. The types of waves that are useful for understanding Earth’s interior are called body waves, meaning that, unlike the surface waves on the ocean, they are transmitted through Earth materials.

Imagine hitting a large block of strong rock (e.g., granite) with a heavy sledgehammer (Figure \(\PageIndex{1}\)). At the point where the hammer strikes it, a small part of the rock will be compressed by a fraction of a millimetre. That compression will transfer to the neighbouring part of the rock, and so on through to the far side of the rock—all in a fraction of a second. This is known as a compression wave, and it can be illustrated by holding a loose spring (like a Slinky) that is attached to something (or someone) at the other end. If you give it a sharp push so the coils are compressed, the compression propagates (travels) along the length of the spring and back (Figure \(\PageIndex{2}\)). You can think of a compression wave as a “push” wave—it’s called a P wave (although the “P” stands for “primary” because P waves arrive first at seismic stations).

When we hit a rock with a hammer, we also create a different type of body wave, one that is characterized by back-and-forth vibrations (as opposed to compressions). This is known as a shear wave (S wave, where the “S” stands for “secondary”), and an analogy would be what happens when you flick a length of rope with an up-and-down motion. As shown in Figure \(\PageIndex{2}\), a wave will form in the rope, which will travel to the end of the rope and back.

Compression waves and shear waves travel very quickly through geological materials. As shown in Figure \(\PageIndex{3}\), typical P wave velocities are between 0.5 kilometers per second (km/s) and 2.5 km/s in unconsolidated sediments, and between 3.0 km/s and 6.5 km/s in solid crustal rocks. Of the common rocks of the crust, velocities are greatest in basalt and granite. S waves are slower than P waves, with velocities between 0.1 km/s and 0.8 km/s in soft sediments, and between 1.5 km/s and 3.8 km/s in solid rocks.

Exercise 9.1 How soon will seismic waves get here?

Imagine that a strong earthquake takes place on Vancouver Island within Strathcona Park (west of Courtenay). Assuming that the crustal average P wave velocity is 5 km per second, how long will it take (in seconds) for the first seismic waves (P waves) to reach you in the following places (distances from the epicentre are shown)?

Nanaimo (120 km away)
Surrey (200 km away)
Kamloops (390 km away)

See Appendix 3 for Exercise 9.1 answers.

Mantle rock is generally denser and stronger than crustal rock and both P- and S-waves travel faster through the mantle than they do through the crust. Moreover, seismic-wave velocities are related to how tightly compressed a rock is, and the level of compression increases dramatically with depth. Finally, seismic waves are affected by the phase state of rock. They are slowed if there is any degree of melting in the rock. If the material is completely liquid, P waves are slowed dramatically and S waves are stopped altogether.

As shown on the right-hand part of Figure \(\PageIndex{4}\), the upper approximately 100 km of the Earth is known as the lithosphere. This includes the rigid upper part of the mantle (or lithospheric mantle) and the crust. The next 150 km is the asthenosphere or low velocity zone (because seismic waves are slowed as they pass through that material). As we’ll see below, that part of the mantle is close to it’s melting point and in some regions may be partially molten.

Accurate seismometers have been used for earthquake studies since the late 1800s, and systematic use of seismic data to understand Earth’s interior started in the early 1900s. The rate of change of seismic waves with depth in Earth (as shown in Figure \(\PageIndex{4}\)) has been determined over the past several decades by analyzing seismic signals from large earthquakes at seismic stations around the world. Small differences in arrival time of signals at different locations have been interpreted to show that:

Velocities are greater in mantle rock than in the crust.
Velocities generally increase with pressure, and therefore with depth.
Velocities slow in the area between a 100 and 250 kilometre depth (called the “low-velocity zone”; equivalent to the asthenosphere).
Velocities increase dramatically at 660 kilometre depth (because of a mineralogical transition).
Velocities slow in the region just above the core-mantle boundary (the D” (d-double-prime) layer or “ultra-low-velocity zone”).
S waves do not pass through the outer part of the core.
P wave velocities increase dramatically at the boundary between the liquid outer core and the solid inner core.

One of the first discoveries about Earth’s interior made through seismology was in 1909 when Croatian seismologist Andrija Mohorovičić (pronounced Moho-ro-vi-chich) realized that at certain distances from an earthquake, two separate sets of seismic waves arrived at a seismic station within a few seconds of each other. He reasoned that the waves that went down into the mantle, traveled through the mantle, and then were bent upward back into the crust, reached the seismic station first because although they had farther to go, they traveled faster through mantle rock (as shown in Figure \(\PageIndex{5}\)). The boundary between the crust and the mantle is known as the Mohorovičić discontinuity (or Moho). Its depth is between 30 and 40 kilometers beneath most of the continental crust, and between 5 and 10 kilometers beneath the oceanic crust.

Our current understanding of the patterns of seismic wave transmission through Earth is summarized in Figure \(\PageIndex{6}\). Because of the gradual increase in density (and therefore rock strength) with depth, all waves are refracted (toward the lower density material) as they travel through homogenous parts of Earth and thus tend to curve outward toward the surface. Waves are also refracted at boundaries within Earth, such as at the Moho, at the core-mantle boundary (CMB), and at the outer-core/inner-core boundary.

S waves do not travel through liquids—they are stopped at the CMB—and there is an S wave shadow on the side of Earth opposite a seismic source. The angular distance from the seismic source to the shadow zone is 103° on either side, so the total angular distance of the shadow zone is 154°. We can use this information to infer the depth to the CMB.

P waves do travel through liquids, so they can make it through the liquid part of the core. Because of the refraction that takes place at the CMB, waves that travel through the core are bent away from the surface, and this creates a P wave shadow zone on either side, from 103° to 150°. This information can be used to discover the differences between the inner and outer parts of the core.

Exercise 9.2 Liquid Cores in Other Planets

We know that other planets must have (or at least did have) liquid cores like ours, and we could use seismic data to find out how big they are. The S wave shadow zones on planets A and B are shown. Using the same method used for Earth (on the left), sketch in the outlines of the cores for these two other planets.

Using data from many seismometers and hundreds of earthquakes, it is possible to create a two- or three-dimensional image of the seismic properties of part of the mantle. This technique is known as seismic tomography, and an example of the result is shown in Figure \(\PageIndex{8}\).

The Pacific Plate subducts beneath Tonga and appears in Figure \(\PageIndex{8}\) as a 100 kilometre thick slab of cold (blue-colored) oceanic crust that has pushed down into the surrounding hot mantle. The cold rock is more rigid than the surrounding hot mantle rock, so it is characterized by slightly faster seismic velocities. There is volcanism in the Lau spreading centre and also in the Fiji area, and the warm rock in these areas has slower seismic velocities (yellow and red colors).

Earth's Magnetic Field

EARTH’S MAGNETIC FIELD ORIGINATES IN THE CORE

The liquid outer core is the source of the earth’s magnetic field, as a result of its metallic nature, which means it contains electrons not attached to particular nuclei. Heat is transferred upward to the mantle from the inner core via convective cells, in which the liquid in the outer core flows in looping patterns. The combination of the loose electrons and looping convective flow with the rotation of the earth results in a geodynamo that produces a magnetic field. Because the magnetic field is generated by a dynamically convecting and rotating sphere of liquid, it is unstable. Every now and then, after several hundred thousand to several million years, the earth’s magnetic field becomes unstable to the point that it temporarily shuts down. When it restarts, its north and south magnetic poles must inevitably be reversed, according to the physics of magnetic fields produced spontaneously from geodyamos. (For comparison, the magnetic field of the Sun, which is also produces by convecting electrical charges in a rotating sphere, becomes magnetically unstable and reverses its magnetic field on a more regular basis, every 11 years.)

Given that the inner core is a solid metallic sphere, made mostly of iron and nickel, surrounded entirely by liquid, it can be pictured as a giant ball bearing spinning in a pressurized fluid. Detailed studies of earthquake waves passing though the inner core have found evidence that it is spinning – rotating – just slightly faster than the rest of the earth.

Heat is also being transferred from the solid inner core to the liquid outer core, and this leads to convection of the liquid iron of the outer core. Because iron is a metal and conducts electricity (even when molten), its motion generates a magnetic field.Earth’s magnetic field is defined by the North and South Poles that align generally with the axis of rotation (Figure \(\PageIndex{1}\)). The lines of magnetic force flow into Earth in the northern hemisphere and out of Earth in the southern hemisphere. Because of the shape of the field lines, the magnetic force trends at different angles to the surface in different locations (red arrows of Figure \(\PageIndex{1}\)). At the North and South magnetic poles, the force is vertical. Anywhere near to the equator the force is horizontal, and everywhere in between, the magnetic force is at some intermediate angle to the surface. As we’ll see in Chapter 10, the variations in these orientations provide a critical piece of evidence to the understanding of continental drift as an aspect of plate tectonics.

Earth’s magnetic field is generated within the outer core by the convective movement of liquid iron, but as we discovered in Chapter 8, the magnetic field is not stable over geological time. For reasons that are not completely understood, the magnetic field decays periodically and then becomes re-established. When it does re-establish, it may be oriented the way it was before the decay, or it may be oriented with the reversed polarity. Over the past 250 Ma, there have been a few hundred magnetic field reversals, and their timing has been anything but regular. The shortest ones that geologists have been able to define lasted only a few thousand years, and the longest one was more than 30 million years, during the Cretaceous (Figure \(\PageIndex{2}\)).

Exercise 9.3 What would a magnetic dip meter tell you?

Regular compasses point only to the north magnetic pole, but if you have a magnetic dip meter you could also measure the angle of the magnetic field at your location in the up-and-down sense.

Using Figure \(\PageIndex{1}\) as a guide, describe where you’d be on Earth if the vertical angles are as follows:

Up at a shallow angle
Parallel to the ground
Down at a steep angle
Straight down

See Appendix 3 for Exercise 9.3 answers.

Figure \(\PageIndex{2}\) Magnetic field reversal chronology for the past 170 Ma. The first 5 Ma of the magnetic chronology are shown in more detail in Figure \(\PageIndex{3}\), although the time scale is in the opposite direction in that figure. “Geomagnetic polarity 0-169 Ma” by Anomie. Adapted by Steven Earle. Public domain.

Changes in Earth’s magnetic field have been studied using a mathematical model, and reversals have been shown to take place when the model was run to simulate a period of several hundred thousand years. The fact that field reversals took place shows that the model is a reasonably accurate representation of the Earth. According to the lead author of the study, Gary Glatzmaier, of University of California at Santa Cruz: “Our solution shows how convection in the fluid outer core is continually trying to reverse the field but that the solid inner core inhibits magnetic reversals because the field in the inner core can only change on the much longer time scale of diffusion. Only once in many attempts is a reversal successful, which is probably the reason why the times between reversals of the Earth’s field are long and randomly distributed.” A depiction of Earth’s magnetic field lines during a stable period and during a reversal is shown in Figure \(\PageIndex{3}\). To read more about these phenomena see Glatzmaier’s Geodynamo website.

Figure \(\PageIndex{3}\) Depiction of Earth’s magnetic field between reversals (left) and during a reversal (right). The lines represent magnetic field lines: blue where the field points toward Earth’s centre and yellow where it points away. The rotation axis of Earth is vertical, and the outline of the core is shown as a dashed white circle. The earth’s magnetic fields is normally very uniform with the magnetic field pointing towards the earth in the north and away from the earth in the south. During a reversal, the Earth’s magnetic field becomes very convoluted. “Supercomputer models of Earth’s magnetic field” by NASA. Public domain.

Magnetic Field and Magnetosphere

We can find additional clues about Earth’s interior from its magnetic field. Our planet behaves in some ways as if a giant bar magnet were inside it, aligned approximately with the rotational poles of Earth. This magnetic field is generated by moving material in Earth’s liquid metallic core. As the liquid metal inside Earth circulates, it sets up a circulating electric current. When many charged particles are moving together like that—in the laboratory or on the scale of an entire planet—they produce a magnetic field.

Earth’s magnetic field extends into surrounding space. When a charged particle encounters a magnetic field in space, it becomes trapped in the magnetic zone. Above Earth’s atmosphere, our field is able to trap small quantities of electrons and other atomic particles. This region, called the magnetosphere, is defined as the zone within which Earth’s magnetic field dominates over the weak interplanetary magnetic field that extends outward from the Sun (Figure 8.5).

Illustration of the Earth’s Magnetosphere. At left an arrow points leftward indicating the direction of the Sun. The Solar wind is drawn as numerous particles coming from the left. Slightly off-center to the right the Earth is shown, with an arrow for the north pole pointing upward, and one pointing down for the south pole. To the left and right of the Earth are two nested purple crescents with their points touching the poles of the Earth. These areas are labeled as the Van Allen belts. Outside the Van Allen belts the lines of the magnetic field are drawn in white. On the left side of the Earth (facing the Sun in this diagram), the lines originate at the north pole and curve out away from the surface then curve back to end at the south pole. Four of these curves are shown, each extending further out into space than the one proceeding it. On the right side of the Earth (facing away from the Sun), the magnetic field lines are also drawn, but have very different shapes than those on the left. The innermost line on the right looks very similar to the innermost line of the left. But each successive line moves further and further out into space before returning to the poles. Thus the magnetic field is much more elongated in shape on the side of the Earth facing away from the Sun. — Figure 8.5 : Earth’s Magnetosphere. A cross-sectional view of our magnetosphere (or zone of magnetic influence), as revealed by numerous spacecraft missions. Note how the wind of charged particles from the Sun “blows” the magnetic field outward like a wind sock.

Where do the charged particles trapped in our magnetosphere come from? They flow outward from the hot surface of the Sun; this is called the solar wind. It not only provides particles for Earth’s magnetic field to trap, it also stretches our field in the direction pointing away from the Sun. Typically, Earth’s magnetosphere extends about 60,000 kilometers, or 10 Earth radii, in the direction of the Sun. But, in the direction away from the Sun, the magnetic field can reach as far as the orbit of the Moon, and sometimes farther.

The magnetosphere was discovered in 1958 by instruments on the first US Earth satellite, Explorer 1, which recorded the ions (charged particles) trapped in its inner part. The regions of high-energy ions in the magnetosphere are often called the Van Allen belts in recognition of the University of Iowa professor who built the scientific instrumentation for Explorer 1. Since 1958, hundreds of spacecraft have explored various regions of the magnetosphere. You can read more about its interaction with the Sun in a later chapter.

The Magnetosphere

Image of Earth’s Van Allen Radiation Belts. A cutaway model of the radiation belts with the 2 RBSP satellites flying through them. The radiation belts are two donut-shaped regions encircling Earth, where high-energy particles, mostly electrons and ions, are trapped by Earth’s magnetic field. This radiation is a kind of “weather” in space, analogous to weather on Earth, and can affect the performance and reliability of our technologies, and pose a threat to astronauts and spacecraft. The inner belt extends from about 1000 to 8000 miles above Earth’s equator. The outer belt extends from about 12,000 to 25,000 miles. This graphic also shows other satellites near the region of trapped radiation. — Public Domain | Image courtesy of NASA.

Earth also has a strong Magnetic Field, which is unusual for a Rocky Planet. It is believed to be due to Earth’s molten core undergoing rotation and convection (giving off heat). Mercury and the Moon exhibit very weak magnetic fields, whereas Venus has none and Mars virtually none; yet there is evidence of a magnetic field in Mar’s early history.

The Van Allen Radiation Belts are charged Particle Belts made up of ions and electrons, which were predicted by James Van Allen, and discovered in 1958 by the US Satellite Explorer 1. These belts protect Earth as intense solar particles from the Sun’s solar wind strikes Earth. These solar wind particles can occasionally be seen as Aurora, as the charged solar particles strike Earth’s poles.