Seismology is the study of vibrations within Earth. These vibrations are caused by events such as earthquakes, extraterrestrial impacts, explosions, storm waves hitting the shore, and tidal effects. Seismic techniques have been most widely applied to the detection and study of earthquakes, but seismic waves also provide important information about 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 the Earth.
Imagine hitting a large block of strong rock with a heavy sledgehammer (Figure 3.4). 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, before the rock bounces back — 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 3.5). 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 that has 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 3.4, 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 3.6, typical P-wave velocities are between 0.5 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 3.1 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/s, how long will it take for the first seismic waves (P-waves) to reach you in the following places?
|Location/distance from earthquake||Time (s)|
|Nanaimo (120 km)|
|Surrey (200 km)|
|Kamloops (390 km)|
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 whether 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.
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 change seismic wave velocity with depth in Earth (Figure 3.7) has been determined over the past several decades by analyzing seismic signals from large earthquakes all 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 100 km and 250 km depth (called the “low-velocity zone”; equivalent to the asthenosphere).
- Velocities increase dramatically at 660 km depth (because of a change in minerals).
- Velocities slow in the region just above the core-mantle boundary (the D” 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 made in the early 1900s by Croatian seismologist Andrija Mohorovičić (pronounced Moho-ro-vi-chich). He noticed that sometimes seismic waves arrived at seismic stations further from the earthquake before they arrived at closer ones. He reasoned that the waves that traveled further were faster because they bent down and traveled faster through different rocks (those of the mantle) before being bent upward back into the crust (Figure 3.8). The boundary between the crust and the mantle is known as the Mohorovičić discontinuity (or Moho). Its depth is between 60 km and 80 km beneath major mountain ranges, around 30 km to 50 km beneath most of the continental crust, and between 5 km and 10 km beneath the oceanic crust.
Our current understanding of the patterns of seismic wave transmission through Earth is summarized in Figure 3.9. Because of the gradual increase in density (and therefore rock strength) with depth, all waves are refracted (bent) toward the lower density material as they travel through the 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, so 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 3.2 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 as 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 images of the seismic properties of the mantle. This technique is known as seismic tomography, and an example of the result is shown in Figure 3.10.
The result shows a 100 km thick slab of cold ocean crust (in blue) belonging to the Pacific tectonic plate plunging into the hot mantle beneath Tonga. The cold rock is more rigid than the surrounding hot mantle rock, so seismic waves travel through it more rapidly. There is volcanism in the Lau spreading centre, in the Fiji area, and at the Tonga Arc, and we can see the warm rock in these areas as slower seismic velocities (shown in yellow and red).