Indirect observing techniques are required to probe the internal structures of planets. Here on Earth, the deepest holes geologists have ever drilled only go down ten or twelve kilometers into the crust, which is a mere pinprick in the 6400-kilometer radius of the planet. Despite the famous Jules Verne fantasy, we have no prospect of ever traveling to the center of the Earth. The crushing pressure and increasing temperature of the surrounding rocks would make drilling much farther down than the current record of 12 km extremely difficult. Information about deeper structure cannot come from drill holes. But we can learn about this inaccessible region from the study of seismic waves passing through the Earth. These waves can be generated by earthquakes, volcanism, or artificial explosions.
Just as light deflects when it passes through a prism, seismic waves bend according to the density of the rock they pass through. Just as light slows down when it travels from air into glass, seismic waves slow down when they enter denser rock. After an earthquake, scientists around the world compare the strength and arrival time of the seismic waves (this is why it is so important to solve for the epicenter of earthquakes). They can use that information to build, one earth quake at a time, a 3-dimensional map of the internal structure of the planet. Similar, but more limited, experiments have been carried out on the Moon using seismometers brought by the Apollo astronauts. Since the moon doesn't have as many natural quakes, these seismic detectors are used primarily to detect the vibrations generated by purposely crashing spacecraft onto the moon, and from asteroid impacts.
Seismic studies on both the Earth and Moon provide evidence that the planets, as predicted, increase in density toward their cores. The variation can be explained by the process of differentiation, which concentrated more dense materials at the center of the bodies while they were still molten. Through seismic studies we can map the transitions from dense core, to overlain mantle of moderate density, to a crust of the least dense rocks at the surface.
All the diverse features of our familiar geography – mountains, canyons, cliffs and valleys – are confined to the slender zone of the Earth's crust. The crust ranges from about 5 km thick under the oceans to about 30 km thick on the continents. From the top of the highest mountain to the bottom of the deepest ocean trench is less than 0.1% of the radius of the planet. If constructed on the same scale, the Earth would be as smooth as a billiard ball.
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). Click here for original source URL.
The boundary between the crust and the mantle is marked by a seismic discontinuity called the Mohorovičić (or Moho) discontinuity. This occurs at about 5 to 10 km below the oceanic crust, and about 30 to 60 km below the continents. At this depth, seismic waves suddenly increase their velocity. This is due to an abrupt change in density and in chemical composition from the aluminum-rich rocks of the crust to the magnesium-rich rocks of the mantle. Another discontinuity occurs at a depth of about 70 km, where the mantle transitions from brittle behavior (lithosphere) to plastic behavior (asthenosphere).
We can tell from the density of the core that it can't be pure iron. About 10% has to be other metals, probably nickel, and perhaps even some sulfur. The outer part of the core is a cauldron of molten metal. The inner part is even hotter, but it's under such high pressure that it remains solid, in spite of the heat. Earth's iron core has a radius of about 3500 kilometers, just over half of the planet's radius.
While terrestrial data is by far the best data we have in the solar system, we also have a fair amount of lunar seismology data. The first lunar seismic detectors were installed by Apollo astronauts. The instruments registered some meteorite impacts and a very few moon quakes, showing the Moon to be much less geologically active than the Earth. Using the sparse data, scientists found that the interior of the Moon is essentially the same material as the Earth's mantle, with perhaps a very small iron core. This makes the mean density of the Moon much lower than Earth's. This was suspected even before the seismic experiments, based on a simple consideration of the Moon's orbit. If we know the Earth's mass, the force that keeps the Moon in its orbit tells us the Moon's mass (since the Earth isn't much much bigger than Moon. This doesn't work to determine the mass of most planets going around the Sun.). Divide this mass by the volume of the Moon, and you get the mean density. This density is too low for the Moon to be structured like the Earth, with a large, dense core. The lack of a metal core is an important clue to how the Moon formed, which is discussed in another article.
Diagram of the internal structure of the Earth. Click here for original source URL.