What Causes Earthquakes?
Earthquakes are the release of stored energy within rocks created by the motion of tectonic plates.
Plate tectonics is a result of the cooling of the Earth’s mantle. The Earth produces energy deep within its interior from the decay of radioactive elements. This energy, combined with the heat that was produced during planetary formation, is redistributed to the cooler surface through both deep-seated melting and the slow circulation (convection) of hot rocks beneath the Earth’s crust. Because the hot mantle material is less dense than the material above it, the hot material is more buoyant and rises to the surface. The hotter material eventually loses heat by conduction, advection, and partial melting. Meanwhile, colder, denser rocks sink back into the mantle.
Figure 1. Tectonic plate boundaries.
The mantle convection cells are coupled to the creation and recycling of the crust, which is one of the fundamental processes of plate tectonics. Newly made oceanic crust is formed at spreading centres (i.e., linear features where magma is extruded from the mantle). These spreading centres divide the crust into large rigid lithospheric plates (Figure 1), comprising relatively cool and buoyant crust and underlying mantle.
Figure 2. Downgoing slab on the west coast of Canada shown at left.
The thickness of the oceanic crust varies from as little as 1.6 km to an average of 10–15 km in older lithosphere.
The continental crust varies in thickness from 10 km to almost 70 km in mountainous regions, with an average of 30–40 km. As the new oceanic crust is created, the old crust is shoved aside like a conveyor belt. When an old, dense, and rigid oceanic plate collides with another plate, the denser oceanic crust sinks, or subducts, beneath the more buoyant crust into the mantle (Figure 2). The subducting crust becomes part of a downwelling limb of a mantle convection cell.
Over vast stretches of geological time, oceanic crusts are consumed and continents collide with each other. When plates are consumed or collide, stresses are formed within the crustal elements. The crust is compressed or stretched from these motions, and earthquakes result when a sudden slip releases this stress. The largest earthquakes occur at subduction zones, where an oceanic plate slips beneath a continental plate. A recent example of this was the magnitude 9 Tohoku earthquake 70 km east of the Oshika Peninsula (see Figure 5) on March 11, 2011, which initiated a tsunami.
The stress accumulated in rock masses over time can eventually exceed the mechanical strength of the rocks and fracture them in a planar or sheet-like feature known as a fault. The rupture (or slip) on the fault generates an earthquake. Earthquakes occur on existing faults (whether they are known to us or not) or they create new faults.
The direction of movement on the fault depends on the faults orientation with respect to the stress field and to the presence of fluids within the fault (Figure 3).
Figure 3. Fault blocks showing stress field and planar movement for (a) normal fault in which the hanging wall moves downward relative to the foot wall, (b) strike-slip in which the blocks move laterally in opposing directions , and (c) reverse or thrust fault in which the hanging wall moves upward relative to the foot wall. The three orthogonal stress directions are shown for vertical (SV), maximum horizontal (SHmax) and minimum horizontal (SHmin) stresses
How do Seismologists Record Earthquakes?
Earthquakes result from the sudden slip of rock along a fault. Movement along the fault sends seismic waves through the surrounding rock, some of which travel to the surface of the Earth. An analogy is a pebble thrown into a pond, creating ripples that travel through the water in all directions. The atoms within the rock respond to the shockwave in a similar way to the water in the pond, by moving (Figure 4) back and forth parallel to the direction the wave is travelling (primary or compressional waves “P”) and up and down or side to side in a direction perpendicular to wave propagation (secondary shear waves “S”). Energy is transferred through the rocks, moving away from the focus of the earthquake and, like the ripples in water, decay with distance.
Figure 4a: Particle movement of seismic waves.
Shallow earthquakes can generate other types of seismic waves known as surface waves, which are trapped in the surface layers of the crust. These are trapped shear waves (Love waves) and combinations of shear and compressional waves (Rayleigh waves). It is the shearing and rolling action of surface waves and, to a lesser degree, the secondary waves travelling up through the crust that cause the most damage to structures.
To record earthquakes, seismologists setup a device called a seismometer. Seismometers passively “listen” to the shaking of the ground. When an earthquake occurs, the seismometer records the incoming seismic waves (P, S, and surface waves) that reach the surface.
Figure 4b. Nanometrics Trillium Compact Seismometer and Taurus digitizer/datalogger.
How do we measure the “size” of an Earthquake?
The energy released during an earthquake, quantified by its magnitude (M), is directly proportional to the planar area (size) of the fault and the amount of slippage on the fault; large faults can generate large earthquakes. An early magnitude scale was developed by Charles Richter in 1935. The local (Richter) magnitude (ML) is calculated from measurements of the largest peak of ground displacement shown on the vertical trace of the seismogram. The peak displacement is normalized by the time to complete one cycle, corrected for the distance the wave travelled to the seismometer, and scaled to duplicate the result that would have been recorded on an instrument similar to that used by Richter. Richter designed the units on the magnitude scale so that each unit marked a tenfold increase in size (logarithmic), in order to address the large range of values. Although the Richter scale is popular in the media, it is generally restricted to small earthquakes.
Traditionally, for large shallow earthquakes (5 ML or greater), the surface wave peak displacement was used to calculate the magnitude (MS), and for large deep earthquakes, the displacement of the primary body wave was used to calculate the magnitude (Mb). These magnitudes, however, tend to saturate (reach a maximum value) at values over 8. For this reason, the moment magnitude (MW) is now preferred for earthquakes larger than M3.5 ML and at any depth. The MW is calculated from the seismic moment (MO), which is based on the product of the slippage area, the total displacement of the fault, and the rigidity of the rocks.
In general, the closer you are to the source of an earthquake, the more likely you are to feel the earth move beneath your feet. Richter magnitudes larger than about 3 can be felt by humans, and earthquakes larger than about Richter magnitude 5 in a populated area can cause minor damage to structures. An exception to this is the case when ground conditions alter the discharge of energy. Because seismic waves travel more slowly through unconsolidated sediments than bedrock, when the wave moves from bedrock to sediment, the energy of the seismic wave is used to displace material instead of propagating the seismic wave, causing stronger ground movement.
How is an Earthquake Located?
Seismometers record seismic waves from earthquakes. Their signals are digitally recorded and sent by telemetry to seismologists at processing agencies like the Geological Survey of Canada (GSC) or their counterparts at the United States Geological Survey (USGS). Encoded in the data are several key types of information: the arrival time of the wavefront, the first motion of displacement (up or down, called polarity), the amplitude of displacement, and the duration of displacement. The magnitude of the earthquake is calculated from the amplitude and duration as discussed in the section on magnitudes. The origin time and location (latitude, longitude, and depth) are calculated from the arrival times of seismic waves at three or more stations. For example, how far away an earthquake is can be estimated by the difference in time between P and S wave arrivals, analogous to guessing how far away a lightning strike is by counting the seconds between the flash and the thunder.
Figure 5. Seismic signals from the 9 MW Tohoku earthquake on March 11, 2011, detected by seismic stations in and around Alberta. The red flags mark the arrivals set by the automated detection, association, and location algorithms of Antelope seismic acquisition and database software used at the AGS to analyze events.
To more robustly determine an earthquake location, seismologists use a computer program, with information about the seismic stations and the local or regional Earth structure, to calculate the location in an iterative process. Accurate locations require many stations and a good spatial distribution of the stations. Large earthquakes have more energy, and their signals are picked up by both local and more distant stations, whereas small earthquakes (less than 2 ML) are only picked up on stations within about 50 km.
Although the depth is the most poorly determined variable in an earthquake location, there are several constraints that help seismologists. Intraplate earthquakes do not occur deeper than about 15–30 km, depending on the regional geotherm. This is because while colder materials are more brittle and fracture, warmer materials behave as a plastic and are more likely to flow. Another constraint is the identification of Rayleigh surface waves, which requires the focal depth to be less than five kilometres.
See also the USGS site for additional details.
- Earthquake Monitoring
- All about Earthquakes
- Induced Seismicity
- Alberta’s Earthquakes