inner layer of the earth (The Core)

Earth’s Core
The outer Core The boundary between Earth’s mantle and outer core, called the core–mantle boundary, is significant because of its varied material properties. P wave velocities drop dramatically as
they cross the core–mantle boundary, and S waves do not penetrate the outer core. Because S waves do not pass through liquids, their absence in the outer core indicates its liquid state. The change in density at the core–mantle boundary from 5.6 to 9.9 g/cm3 is even
larger than the air–rock difference observed at Earth’s surface.

Based on our knowledge of the composition of meteorites, geologists expect Earth to contain a significantly higher percentage of iron than is observed in rocks found in the crust and mantle. This fact, coupled with the great density of the core, indicates that the outer core consists mostly of iron, with lesser amounts of nickel, which has a density similar to that of iron.

Density and seismic studies suggest that in addition to iron and nickel, about 15 percent of the outer core consists of lighter elements. Based on mineral physics experiments, these are likely to include sulfur, oxygen, silicon, and hydrogen. For instance, an iron
sulfur mixture melts at a much lower temperature than pure iron. As Earth was forming, high-velocity impacts of nebular debris and the decay of radioactive elements caused the temperature of our planet to increase  When heated sufficiently, iron in the presence of sulfur would begin to melt before pure iron. Melting produced liquid blobs of an iron–sulfur alloy that sank to form the core. In a similar manner, other light elements were dragged down into the                                                      core.
The core accounts for about one-sixth of Earth’s volume but one-third of its mass because it is composed mostly of iron, which is the most dense of the common elements. In fact, iron is Earth’s most abundant element when measured by mass.
The Inner Core:
At the center of the core is the inner core, a solid dense sphere (around 13 g/cm3) of iron with trace amounts of other elements. Because the inner core is a sphere whereas Earth’s other layers
are shells, drawings make the inner core appear much larger than it really is (see Fig1).

The inner core is actually relatively small, only 1/142 of the volume of Earth (less than 1 percent). The inner core did not exist
early in Earth’s history, when our planet was hotter. However, as Earth cooled, iron began to crystallize at the center to form the solid inner core. Even today, the inner core continues to grow larger at the expense of the outer core—as the planet cools. Separated from the mantle by the liquid outer core, the solid inner core is free to rotate independently from Earth’s outer layers. Recent studies suggest that the inner core rotates faster than the crust and mantle, lapping them every few hundred years
(Fig12).

 The inner core’s small size and great distance from the surface make it the most difficult-toexamine region of Earth. Discovering Boundaries: The Inner Core–outer Core Boundary The boundary between the solid inner core and liquid outer core was discovered in 1936 by Danish seismologist Inge Lehman. By examining seismograph records, Lehman discovered that some P waves were strongly refracted (bent) by a sudden increase in seismic velocities at a boundary within Earth’s core. The waves she observed were bent enough to arrive within the P-wave shadow zone, as shown in Fig11A.

Because of the work of Lehman and other scientists, it is now understood that the inner core is an integral part of the activities of Earth’s interior.

Middle layer of earth (Mantle)

More than 82 percent of Earth’s volume is contained within the mantle, a nearly 2900-kilometer-thick shell extending from the base of the crust (Moho) to the liquid outer core (see Figure 1). Because S waves readily travel through the mantle, we know that it is a solid rocky layer composed of silicate minerals that are rich in iron and magnesium. However, despite its solid nature, rock in the mantle is quite hot and capable of flow, albeit at very slow velocities.

figure 1 show's that Earth’s layered structure: The properties of Earth’s layers include the physical state of the material (solid, liquid, or gas) as well as

how strong the material is—for example, the distinction between the strong lithosphere and weak asthenosphere. Studies have shown that Earth’s layers are mainly determined by density, with the heaviest materials (iron) at the center and the lightest ones (gases and liquids) on the outside.
The upper Mantle:
Earth’s upper mantle extends from the Moho to a depth of about 660 kilometers and can be divided into three shells:
1. The uppermost mantle is called the lithospheric mantle, and it ranges in thickness from only a few kilometers under the mid-oceanic ridges to perhaps as much as 200 kilometers under the stable continental interiors. This layer and the crust make up Earth’s rigid outer shell, called the lithosphere.
2. Beneath the lithospheric mantle is a weak layer called the asthenosphere. The lithospheric mantle and asthenosphere are compositionally similar; however, lithospheric mantle is strong, and the asthenosphere is weak, as a result of Earth’s temperature structure.
3. The lower portion of the upper mantle, at depths between 410 and 660 kilometers, is called the transition zone.
Rocks brought to the surface by volcanism and other geologic processes have provided geologists with valuable information about the composition of the upper mantle, which is composed mainly of the rock peridotite
(Fig 9).

Peridotite is an ultramafic rock that consists of the minerals olivine and pyroxene, minerals that are rich in iron and magnesium. As a result, the mantle is denser than either the continental crust or the oceanic crust that lie above it. At the depth (and pressures) of the transition zone, the mineral olivine, which is stable in the uppermost mantle, is subjected to greater pressure and collapses into denser structures. In the top half of the transition zone, olivine converts to a more compact structure similar to the mineral spinel and pyroxene converts to a garnet-like structure. Water cycles slowly through Earth and is brought into the mantle by subducting oceanic lithosphere and  carried back to the surface by rising plumes of mantle rock. Mineral physics experiments have revealed that the transition zone is capable of holding a great deal of water up to 2 percent of its weight. This is considerably more than for the rocks of the uppermost mantle, which can hold only about 0.1 percent of their weight as water. Because the transition zone represents 10 percent of Earth’s volume, it could potentially hold up to five times the volume of Earth’s oceans. How much water is actually contained within the transition zone has not been determined.
The Lower Mantle:
The lower mantle lies between the transition zone (660 kilometers) and the liquid core (2900 kilometers). Beneath the 660-kilometer
discontinuity, both olivine and pyroxene take the  structure of the mineral perovskite (Fe, Mg) SiO3 and related minerals. Because the lower mantle is undoubtedly Earth’s largest layer, occupying 56 percent of the volume of the planet, perovskitestructured silicate minerals are the single most abundant material within Earth.
The D" Layer:
In the lowest few hundred kilometers of the mantle is a highly
variable and unusual region called D" (pronounced “dee double-prime”). The D" layer, the boundary layer between the rocky mantle and the liquid iron outer core, is thought to have large variations in composition as well as temperature (Fig10)

Cool areas in the D" layer are thought to be the graveyard of subducted oceanic lithosphere, whereas the hot areas are the birthplace of deep mantle plumes. The very base of D", the part of the mantle in direct contact with the hot liquid iron core, is like
Earth’s surface in that there are “upside-down mountains” of rock that protrude into the core. Furthermore, in some regions of the core–mantle boundary, the base of D" may be hot enough to be partially molten. Evidence for partial melting comes from zones at the very base of the mantle where S-wave velocities decrease by 30 percent, an indication that the material there is quite weak. 

Discovering Boundaries: The Core–Mantle Boundary Evidence that Earth has a distinct central core was uncovered in 1906 by British geologist Richard Dixon Oldham. At locations beyond approximately 100 degrees from the epicenter of a largeearthquake, Oldham observed that P and S waves were absent or very weak. In other words, Oldham found evidence for a central core that produced a shadow zone for seismic waves (Fig 11)

. In 1914,Beno Gutenberg calculated 2900 kilometers to the core boundary depth, which remains the accepted value.
As Oldham predicted, Earth’s core exhibits markedly different properties from the mantle above, which causes considerable refraction of P waves—similarly to how light is refracted as it passes from air to water. In addition, because the outer core is liquid iron, it blocks the transmission of S waves. (Recall that S waves do not travel through liquids.)
The locations of the P- and S-wave shadow zones and how their paths are affected by the core are shown in Fig 11. Whereas some P and S waves still arrive in the shadow zone, they differ greatly from those expected in a planet without a core.

Earth’s Layers (1st upper layer crust)

Combining the data obtained from seismological studies and mineral physics experiments has given us a layer-by-layer understanding of Earth’s composition. The variations in seismic velocities with depth are shown in Fig 7

By examining the behavior of a variety of rocks at the pressures corresponding to these depths, geologists have made important discoveries about the compositions of Earth’s crust, mantle, and core.

Earth’s Crust: (1st layer of earth)

Earth’s crust consists of two distinct types—continental crust and oceanic crust. Continental crust and oceanic crust have very different compositions, histories, and ages. In fact, oceanic crust is compositionally more similar to the mantle than to the continental crust.
Oceanic Crust:
The ocean crust averages about 7 kilometers (4.5 miles) thick and forms at mid-ocean ridges, which separate two diverging tectonic plates. Ocean crust has a density of about 3.0 g/cm3, which compares to measured values for the rocks basalt and gabbro.
Continental Crust:
While oceanic crust is fairly uniform, no two continental regions have the same structure or composition. Continental crust averages about 40 kilometers (25 miles) thick but can be more than 70 kilometers (45 miles) thick in mountainous regions such as the
Himalayas and the Andes. The thinnest crust in North America is beneath the Basin and Range region in the western United States, where the crust is as thin as 20 kilometers (12 miles). The thickest North American crust, beneath the Rockies, is more than 50 kilometers (30 miles) thick. Seismic velocities within continents are quite variable, suggesting that the composition of continental crust must also vary greatly. Continents have an average density of about 2.7 g/cm3, which is much lower than the densities of both oceanic crust and mantle rock. This low density explains why continents are buoyant—acting like giant rafts, floating atop tectonic plates—and why they cannot be readily subducted into the mantle.

Discovering Boundaries: The Moho
The boundarybetween the crust and mantle, called the Moho, was one of the first features of Earth’s interior discovered using seismic waves. In 1909, Croatian seismologist Andrija Mohorovicic discovered this boundary that now bears his name. At the base of the continents, P waves travel about 6 kilometers per second (km/s) but abruptly increase to 8 km/s at a slightly greater depth.Mohorovicic cleverly used this large jump in seismic velocity to make his discovery. He noticed that two different sets of seismic waves were recorded at seismographs located within a few hundred kilometers of an earthquake. One set of waves moved through the ground at about 6 km/s, while the other set of waves traveled about
8 km/s—allowing Mohorovicic to correctly determine that the different waves were traveling through two different layers.During a shallow earthquake, direct waves travel along a nearly straight path through the crust, as shown in Fig 8.

Other seismic waves follow a path through the crust and along the top of the mantle. These are called refracted waves because they are bent, or refracted, as they enter the mantle. Seismographs near the
epicenter record the direct waves first. However, seismographs further from the epicenter record the refracted waves first. The point at which both waves arrive at the same time, called the cross-over, can be used to determine the depth of the Moho. Thus, using data from these two sets of waves and seismographs at various distances from an earthquake’s epicenter, the thickness of the crust
for any location can be calculated. The difference between travel times for direct and refracted waves is comparable to driving to a destination on local roads versus on interstate highways. For short
distances, you will typically arrive sooner if you drive the most direct route using local roads. For long distances, the trip may take less time if you take a less direct route that involves mostly interstate highways. The cross-over point, where both routes take an equal amount of time, is directly related to how far you must drive before reaching the interstate highway. Applied to determining the depth of the Moho, the cross-over is related to how far seismic waves travel through the crust (slow layer) before they reach the mantle (fast layer): The greater the cross-over distance, the deeper the Moho. The Moho lies about 25 to 70 kilometers (15 to 45 miles) beneath the continents and about 5 to 10 kilometers (3 to 6 miles) below the ocean floor.

Describe how seismic waves are used to probe Earth’s interior

Probing earth’s interior

Discovering the structure and properties of Earth’s deep interior has not been easy. Light does not
travel through rock,so we must find another ways to “see” into our  planet earth . The most accurate way to learn about Earth’s interior would be to dig or drill a hole and examine what is extracted. Unfortunately, this is only possible at shallow depths. The deepest a drill has ever penetrated is only 12.3 kilometers (7.5 miles), about 1/500 of the way to Earth’s centre—an extraordinary accomplishment due to the rapid increases in temp and pressure with depth we are going beneath the earth.

“Seeing” Seismic Waves
Fortunately for seismologists, many earthquakes are large enough that their seismic waves travel all the way through Earth and can be detected on the other side (Figure 2).

This property of seismic waves is similar to how medical X-rays image our bones and organs. There are about 100 to 200 earthquakes each year that are large enough (about M 6) to be recorded by seismographs around the globe. These large quakes provide the means to “see” into our planet and have been the source of much of the data that allow us to more fully understand Earth’s interior.

Seismic Velocities

Recall that the various seismic waves travel at different speeds. In addition, the speed at which P waves and S waves travel through Earth’s layers depends largely on the properties of the materials that transmit them. In general, seismic waves travel fastest when rock is stiff (rigid) or less compressible. These properties of stiffness and compressibility are used to interpret the composition and temperature of the rock.
For instance, when rock is heated, it becomes less stiff (imagine warming a frozen chocolate bar), and waves travel through it more slowly. When P waves enter the outer core, which is liquid, they slow dramatically while S waves are not transmitted  Likewise, waves travel at different speeds through Earth materials that have different compositions. For example, seismic waves travel faster through oceanic crust, which is composed of basalt, than through the continental crust, which has an overall composition akin to granite. Thus, the speed at which seismic waves travel through a layer can help determine both the type of material and its temperature.

Interactions between Seismic Waves and Earth’s Layers:

Interpreting the waves recorded on seismograms is challenging because seismic waves usually do not travel along straight paths. Instead, seismic waves interact with Earth’s layers and are reflected and refracted as they pass through our planet (Fig 3).

 You are familiar with reflected sound waves that we call echoes. When a seismic wave hits a boundary between different Earth materials, such as the boundary between the crust and the mantle, some of the waves are reflected back toward the surface (Fig 4).

The remaining energy passes though the boundary and is refracted (bent). This is similar to how light is refracted (bent) as it passes from air to water.
One of the most noticeable behaviour  of seismic waves is that they follow strongly curved (refracted) paths because their velocities generally increase with depth (Fig 5).

 Within a particular layer, the speed of seismic waves increases with depth because pressure increases and squeezes the rock into a more compact, rigid material. Within Earth’s mantle, where there are both distinct boundaries and gradual seismic velocity changes caused by changes in mineral properties, the pattern of seismic waves is complex.

Fig 6

shows what S waves look like when they travel from a deep earthquake through the mantle. Note how the single wave from the shock is soon broken into many different waves that appear on seismograms as separate signals.