- Stress and strain
- Elastic rebound theory - how earthquakes begin
- Seismic waves
- Locating earthquake epicenters
- Determining earthquake depth
- Calculating earthquake magnitude
- Ranking earthquake intensity
- Earthquakes and plate boundaries
- Intraplate earthquakes
- Earthquakes and volcanoes
- Earthquake hazards
- Ground shaking
- Permanent ground displacement
- Earth rupture
- Landslides and avalanches
- Mitigating earthquake damage
- Seismicity and earthquake prediction
Earthquakes are caused by the abrupt release of energy in the earth. The energy moves outward from its source in the form of seismic waves, which cause the earth's surface to shake, making an earthquake.
Most earthquakes are caused by sudden slippage of sections of the crust along faults. This sudden slippage is often referred to as failure of the fault. Other causes of earthquakes include magma movement in the crust, volcanic eruptions, abrupt reduction in the volume of minerals in a subducting plate as they adjust to higher pressures in the mantle, bombs set off by humans, and meteorite impacts. Earthquakes are a response of earth's interior to forces that push, pull, or squeeze its rocks.
Rocks are also subjected to the three types of directed (non-uniform) stress - tension, compression, and shear.
- Tension is a directed (non-uniform) stress that pulls rock apart in opposite directions. The tensional (also called extensional) forces pull away from each other.
- Compression is a directed (non-uniform) stress that pushes rocks together. The compressional forces push towards each other.
- Shear is a directed (non-uniform) stress that pushes one side of a body of rock in one direction, and the opposite side of the body of rock in the opposite direction. The shear forces are pushing in opposite ways.
In response to stress, rock may undergo three different types of strain - elastic strain, ductile strain, or fracture.
- Elastic strain is reversible. Rock that has undergone only elastic strain will go back to its original shape if the stress is released.
- Ductile strain is irreversible. A rock that has undergone ductile strain will remain deformed even if the stress stops. Another term for ductile strain is plastic deformation.
- Fracture is also called rupture. A rock that has ruptured has abruptly broken into distinct pieces. If the pieces are offset - shifted in opposite directions from each other - the fracture is a fault.
Earth's rocks are composed of a variety of minerals and exist in a variety of conditions. In different situations, rocks may act either as ductile materials that are able to undergo an extensive amount of ductile strain in response to stress, or as brittle materials, which will only undergo a little or no ductile strain before they fracture. The factors that determine whether a rock is ductile or brittle include:
- Composition - Some minerals, such as quartz, tend to be brittle and are thus more likely to break under stress. Other minerals, such as calcite, clay, and mica, tend to be ductile and can undergo much plastic deformation. In addition, the presence of water in rock tends to make it more ductile and less brittle.
- Temperature - Rocks become softer - more ductile - at higher temperature. Rocks at mantle and core temperatures are ductile and will not fracture under the stresses that occur deep within the earth. The crust, and to some extent the lithosphere, are cold enough to fracture if the stress is high enough.
- Lithostatic pressure - The deeper in the earth a rock is, the higher the lithostatic pressure it is subjected to. High lithostatic pressure reduces the possibility of fracture because the high pressure closes fractures before they can form or spread. The high lithostatic pressures of the earth's sub-lithospheric mantle and solid inner core, along with the high temperatures, are why there are no earthquakes deep in the earth.
- Strain rate - The faster a rock is being strained, the greater its chance of fracturing. Even brittle rocks and minerals, such as quartz, or a layer of cold basalt at the earth's surface, can undergo ductile deformation if the strain rate is slow enough.
Most earthquakes occur in the earth's crust. A smaller number of earthquakes occur in the uppermost mantle (to about 700 km deep) where subduction is taking place. Rocks in the deeper parts of the earth do not undergo fracturing and do not produce earthquakes because the temperatures and pressures there are high enough to make all strain ductile. No earthquakes originate from below the the earth's upper mantle.
The following correlations can be made between types of stress in the earth, and the type of fault that is likely to result:
- Tension leads to normal faults.
- Compression leads to reverse or thrust faults.
- Horizontal shear leads to strike-slip faults.
Correlations between type of stress and type of fault can have exceptions. For example, zones of horizontal stress will likely have strike-slip faults as the predominant fault type. However there may be active normal and thrust faults in such zones as well, particularly where there are bends or gaps in the major strike-slip faults.
To give another example, in a region of compression stress in the crust, where sheets of rock are stacked on active thrust faults, strike-slip faults commonly connect some of the thrust faults together.
Active faults are faults that have undergone offset, probably during earthquakes, during recent geologic time, in other words, during the Holocene epoch, the last 12,000 years or so. Active faults are considered likely to produce more earthquakes in the future. Ancient faults, which have not undergone offset in recent geologic time and are no longer active, can be found in any extensive exposure of ancient bedrock. Most faults in the earth's crust are ancient, inactive faults, no longer liable to produce any earthquakes. In the United States, the Geological Survey keeps maps and a data base that compile the known active faults in the country. However, there are many active faults that have not been identified, usually because they are hidden beneath surface sediments and vegetation and have not undergone failure in the last few hundred years.
Where a fault has undergone vertical offset that has ruptured the earth's surface, a fault scarp, an exposure of the fracture offset on the earth's surface, is formed. However, erosion and sedimentation processes can erase fault scarps, sometimes within a few years. Faults that have undergone horizontal offset may not produce a scarp, but will produce linear features such as fault trenches that can fill with water and become elongate ponds, offsets in stream drainages, or elongate ridges.
Some faults do not fail all the way to the earth's surface. Faults that remain hidden underground, without any signs of rupture, escarpment, or lineation on the earth's surface are called blind faults.
Prior to the the great earthquake that destroyed most of San Francisco in 1906, it was thought that earthquakes caused the earth to fracture, forming faults. In the days following that earthquake, geologist Harry Fielding Reid observed and measured the extent of rupture and offset of the San Andreas fault that had occurred during the earthquake. As a result of his study, he published a paper in which he proposed that it was the fracture of the fault that caused the earthquake, rather than the other way around. His theory of how earthquakes occur as a result of fracture of rocks along a fault is called elastic rebound theory.
According to elastic rebound theory, the blocks of rock on opposite side of a fault are forced in opposite directions by stress. The masses of rock undergo elastic strain, slowly flexing and moving in opposite directions. Eventually the internal strength of the rock along the fault is reached and brittle failure - rupture - occurs, producing an earthquake. The bodies of rock on either side of the fault "rebound," returning to their original shape, now offset from the other block of rock along the fault where the rock has ruptured and shifted.
Elastic rebound theory has been verified by precise GPS measurements of displacements of rock around and along a fault before, during, and after an earthquake. Although elastic rebound theory is thought to be correct in general, there are differences in the details of how each fault breaks and how each earthquake is produced by fault movement. These details differ from fault to fault and earthquake to earthquake.
The point on a fault within earth's crust where the fracturing begins and most slippage occurs is called the focus of the earthquake. Another name for it is the hypocenter. The point on the earth's surface directly above the focus is the epicenter. The epicenter is not where the earthquake originated. Earthquakes originate within the earth. The epicenter is the point on the surface of the earth directly above where the earthquake originated.
When an earthquake occurs, some of the energy it releases is turned into heat within the earth. Some of the energy is expended in breaking and permanently deforming the rocks and minerals along the fault. The rest of the energy, which is most of the energy, is radiated from the focus of the earthquake in the form of seismic waves.
Seismic waves fall into two general categories: body waves, which travel through the interior of the earth, and surface waves, which travel only at the earth's surface.
There are two types of body waves: P-waves and S-waves. The P in P-waves stands for primary, because these are the fastest seismic waves and are the first to be detected once an earthquake has occurred. P-waves travel through the earth's interior many times faster than the speed of a jet airplane, taking only a few minutes to travel across the earth.
P-waves are predominantly compressional waves. As a P-wave passes, material compresses in the same direction the wave is moving, and then extends back to its original thickness once the wave has passed. The speed at which P-waves travel through material is determined by:
- rigidity -- how strongly the material resists being bent sideways and is able to straighten itself out once the shearing force has passed - the more rigid the material, the faster the P-waves
- compressibility -- how much the material can be compressed into a smaller volume and then recover its previous volume once the compressing force has passed - the more compressible the material, the faster the P-waves
- density -- how much mass the material contains in a unit of volume - the greater the density of the material, the slower the P-waves
The animations below show P-waves propogating across a plane (left) and from a point source (right). They are from wikipedia.org/wiki/P-wave uploaded November, 2006 by Christophe Dang Ngoc Chan.
P-waves travel through liquids and gases as well as through solids. Although liquids and gases have zero rigidity, they have compressibility, which enables them to transmit P-waves. Sound waves are P-waves moving through the air.
Because the earth's mantle becomes more rigid and compressible as the depth below the asthenosphere increases, P-waves travel faster as they go deeper in the mantle. The density of the mantle also increases with depth below the asthenosphere. The higher density reduces the speed of seismic waves. However, the effects of increased rigidity and compressibility in the deep mantle are much greater than the effect of the increased density.
|P-waves travel through materials with rigidity and/or compressiblity, and density|
|greater rigidity||faster P-waves|
|greater compressibility||faster P-waves|
|greater density||slower P-waves|
The S in S-waves stands for secondary, because they are the second-fastest seismic waves and the second type to be detected once an earthquake has occurred. Although S-waves are slower than P-waves, they still travel fast, over half the speed of P-waves, moving at thousands of kilometers per hour through the earth's crust and mantle.
S-waves are shear waves (though that is not what the S stands for). They move by material flexing or deforming sideways (shearing) from the direction of wave travel, and then returning to the original shape once the wave passes. The speed at which S-waves travel through material is determined only by:
- rigidity -- how strongly the material resists being bent sideways and is able to straighten itself out once the shearing force has passed - the more rigid the material, the faster the S-waves
- density -- how much mass the material contains in a unit of volume - the greater the density of the material, the slower the S-waves
The animations below show S-waves propogating across a plane (left) and from a point source (right). They are from wikipedia.org/wiki/S-wave uploaded November, 2006 by Christophe Dang Ngoc Chan.
S-waves can travel only through solids, because only solids have rigidity. S-waves cannot travel through liquids or gases.
Because the earth's mantle becomes more rigid as its depth below the asthenosphere increases, S-waves travel faster as they go deeper in the mantle. The density of the mantle also increases at greater depth, which has the effect of reducing the speed of seismic waves, but the increase in rigidity is much greater than the increase in density, so S-waves speed up as they get deeper in the mantle, in spite of the increased density.
|S-waves travel through materials with rigidity and density|
|greater rigidity||faster S-waves|
|greater density||slower S-waves|
There are two types of surface waves, Rayleigh waves and Love waves. Rayleigh waves are named after Lord Rayleigh (John Strutt), an English aristocrat who, in his work as a scientist and mathematician, developed a detailed mathematical accounting of the type of surface wave named after him. Rayleigh waves are set off by the combined effect of P- and S-waves on the earth's surface. Rayleigh waves are sometimes called rolling waves. In Rayleigh waves the surface of the earth rises up and sinks down in crests and troughs, similar to waves on the surface of water. People who are outdoors during a major earthquake commonly see Rayleigh waves moving across the surface of the earth, and can feel the ground rising and falling as the waves pass beneath them.
Love waves, sometimes called L-waves, are named after Augustus Love, an English mathematician and physicist who first modeled them mathematically. Love waves involve the surface shearing sideways and then returning to its original form as each wave passes.
All surface waves travel slower than body waves and Rayleigh waves are slower than Love waves.
Seismology is the study of seismic waves. Seismology is also the study of earthquakes, mainly through the waves they produce. By measuring and analyzing seismic waves, seismologists can derive such information as:
- The epicenter of an earthquake
- The depth of an earthquake focus
- The magnitude (power) of an earthquake
- The type of fault movement that produced an earthquake
- Whether an earthquake beneath the ocean is likely to have generated a tsunami (a set of giant ocean waves)
In addition to information about earthquakes and faults, seismology gives us knowledge of the layers of the earth. Much of what we know about the crust, lithosphere, asthenosphere, mantle, and core comes from seismology. See the Earth's interior Basics page.
Seismology also gives us information about underground nuclear testing that takes place anywhere on earth, allows possible oil reservoirs to be located within the earth's crust, and helps us predict when a volcano is about to erupt.
Seismographs and seismometers are the instruments used to measure seismic waves. The traditional analog seismograph utilizes a pen (stylus) embedded in a heavy weight, which is suspended on springs. When the earth moves during an earthquake, a piece of paper rolling beneath the stylus moves with the earth, but the stylus, with its weight suspended on springs, remains stationary, drawing lines on the sheet of paper that show the seismic motions of the earth. The USGS photo below shows a seismogram from a seismograph located in Columbia, California that recorded the 1989 Loma Prieta Earthquake.
With modern technology, seismographs with pens and rolling sheets of paper are being replaced by seismometers with electronic sensors and computer screens. Seismographs and seismometers both produce a seismogram, which is a graphic record of the seismic waves, viewed either on paper or on a computer monitor.
Because P-waves travel faster than S-waves, the farther away you are from an earthquake, the greater the time lag between when the first P-wave arrives from the earthquake and when the first S-wave arrives. This difference in arrival time can be used to determine the distance to the earthquake. If three seismograph stations, at three different locations, each determine their distance to the earthquake, the location of the earthquake epicenter can be determined by triangulation. From each seismograph station location on a map, a circle is drawn with a radius equal to the distance to the earthquake from the seismograph station. The earthquake epicenter must be somewhere on that circle. Where all three circles intersect at a single point, that point is the earthquake epicenter.
In today's world, many seismograph stations share data via the Internet. Once an earthquake occurs, data from several seismograph stations, usually more than three, is combined and calculated by computer to determine a statistically best fit epicenter location.
The depth of the earthquake focus is also calculated using combined seismograph data from several stations, but it is a more difficult calculation. As a result, the calculated depth of an earthquake is somewhat less precise than the determination of the epicenter location. For example, an epicenter may be located to within 1 km, but the depth to the focus may be plus or minus several kilometers.
Most earthquakes occur in the crust or upper lithosphere, less than 70 km (50 miles) deep in the earth. Earthquakes that are less than 70 km deep are called shallow earthquakes. Earthquakes that occur between 70 and 300 km deep are intermediate-depth earthquakes. Earthquakes 300-700 km deep are called deep earthquakes. None deeper have been detected.
Most intermediate-depth earthquakes, and all deep earthquakes, occur in subduction zones. In these zones, subducting oceanic lithosphere remains cold and brittle, prone to undergoing fracture and producing earthquakes, even as it sinks down into mantle depths.
In theory, once a subducting plate is several hundred km deep, its temperature should have increased enough by conduction of heat from surrounding mantle, and the lithostatic pressure become high enough, to prevent any brittle behavior. Earthquakes at that depth, therefore, should not occur by fracture and slippage along faults the way most earthquakes do in the shallow crust and upper lithospheric mantle. The deepest earthquakes, down to 700 km, are probably due instead to olivine, the most abundant mineral in the subducting plate, contracting into a new crystal lattice with a denser structure. This contraction occurs in response to the increased lithostatic pressure encountered by the subducting slab as it sinks deep into the mantle.
The magnitude of an earthquake is a number that allows earthquakes to be compared with each other in terms of their relative power. For several decades, earthquake magnitudes were calculated based on a method first developed by Charles Richter, a seismologist based in California. Richter used seismograms of earthquakes that occurred in the San Andreas fault zone to calibrate his magnitude scale.
Two measurements are factored together to determine the Richter magnitude of an earthquake: the amplitude of the largest waves recorded on a seismogram of the earthquake, and the distance to the epicenter of the earthquake. The maximum amplitude seismic wave - the height of the tallest one - is measured in mm on a seismogram. The distance to the epicenter must also be taken into account because the greater the distance from the earthquake, the smaller the waves get. The effect of distance is factored out of the calculation. There is no upper limit defined for the Richter scale, but after a century of seismograph measurements, it appears that rocks in the earth release their stress before building up enough energy to reach magnitude 10.
The Richter scale was found to not transfer very well from the San Andreas fault zone, a transform plate boundary, to the much more powerful earthquakes that occur at convergent plate boundaries, particularly subduction zone earthquakes. Therefore, the Richter scale has been replaced by the moment magnitude scale, symbolized as Mw.
The moment magnitude scale is broadly similar to the Richter scale, but it takes more factors into account, including the total area of the fault that moves during the earthquake, and how much it moves. This produces a magnitude number that is a better indicator of the total amount of energy released by the earthquake. Because the moment magnitude scale has replaced the Richter scale, we will assume from here on that we are referring to moment magnitude, not Richter magnitude, when we speak of earthquake magnitude.
The magnitude scale portrays energy logarithmically to approximately base 32. For example, a magnitude 6.0 earthquake releases about 32 times as much energy as a magnitude 5.0 earthquake. A magnitude 7.0 releases about 32 x 32 = 1024 times as much energy as a magnitude 5.0 earthquake. A magnitude 9.0 earthquake, which rarely occurs, releases over a million times as much energy as a magnitude 5.0 earthquake.
Earthquake intensity is very different from earthquake magnitude. Earthquake intensity is a ranking based on the observed effects of an earthquake in each particular place. Therefore, each earthquake produces a range of intensity values, ranging from highest in the epicenter area to zero at a distance from the epicenter. The most commonly used earthquake intensity scale is the Modified Mercalli earthquake intensity scale. Refer to the Modified Mercalli Intensity Scale page on the US Geological Survey Earthquake Hazards Program website for an abbreviated version.
The table below shows approximately how many earthquakes occur each year in each magnitude range and what the intensity might be at the epicenter for each magnitude range.
|Magnitude||Average number per year||Modified Mercalli Intensity||Description|
|0 - 1.9||>1 million||--||micro - not felt|
|2.0 - 2.9||>1 million||I||minor - rarely felt|
|3.0 - 3.9||about 100,000||II - III||minor - noticed by a few people|
|4.0 - 4.9||about 10,000||IV - V||light - felt by many people, minor damage possible|
|5.0 - 5.9||about 1,000||VI - VII||moderate - felt by most people, possible broken plaster and chimneys|
|6.0 - 6.9||about 130||VII - IX||strong - damage variable depending on building construction and substrate|
|7.0 - 7.9||about 15||IX - X||major - extensive damage, some buildings destroyed|
|8.0 - 8.9||about 1||X - XII||great - extensive damage over broad areas, many buildings destroyed|
|9.0 and above||< 1||XI - XII||great - extensive damage over broad areas, most buildings destroyed|
Most, but not all, earthquakes occur at or near plate boundaries. A great deal of stress is concentrated and a great deal of strain, much of it in the form of rupture of the earth, takes place at locations where two plates diverge, transform, or converge relative to each other.
Tension is the dominant stress at divergent plate boundaries. Normal faults and rift valleys as the predominant earthquake-related structures at divergent plate boundaries. Earthquakes at divergent plate boundaries are usually relatively shallow, and, though they can be damaging, the most powerful earthquakes at divergent plate boundaries are not nearly as powerful as the most powerful earthquakes at convergent plate boundaries.
Transform plate boundaries are zones dominated by horizontal shear, with strike-slip faults the most characteristic fault type. Most transform plate boundaries cut through relatively thin oceanic crust, part of the structure of the ocean floor, and produce relatively shallow earthquakes that are only rarely of major magnitude. However, where transform plate boundaries and their strike-slip faults cut through the thicker crust of islands or the even thicker crust of continents, more stress may need to build up before the thicker masses of rock will rupture, and so the magnitudes of earthquakes can be higher than in transform plate boundary zones confined to thin oceanic crust. This is evident in such places as the San Andreas fault zone of California, where a transform fault cuts through continental crust and earthquakes there sometimes exceed 7.0 in magnitude.
Convergent plate boundaries are dominated by compression. The major faults found in convergent plate boundaries are usually reverse or thrust faults, including a master thrust fault at the boundary between the two plates and typically several more major thrust faults running roughly parallel to the plate boundary. The most powerful earthquakes that have been measured are subduction earthquakes, up to greater than 9.0 in magnitude. All subduction zones in the world are at risk of subduction earthquakes with magnitudes up to or even greater than 9.0 in extreme cases, and are likely to produce tsunamis. This includes the Cascadia subduction zone of northern California and coastal Oregon and Washington, the Aleutian subduction zone of southern Alaska, the Kamchatka subduction zone of Pacific Russia, the Acapulco subduction zone of southern Pacific Mexico, the Central American subduction zone, the Andean subduction zone, the West Indian or Caribbean subduction zone, and subduction zones of Indonesia, Japan, the Phillipines, and several more subduction zones in the western and southwestern Pacific Ocean.
Some earthquakes take place far away from plate boundaries. Earthquakes can occur wherever there is sufficient stress in the earth's crust to drive rocks to rupture.
For example, Hawaii is thousands of km (thousands of miles) from any plate boundary, but the volcanoes that compose the islands have built up so rapidly that they are still undergoing gravitational stabilization. Sectors of the Hawaiian islands occasionally slump along normal faults, producing intraplate earthquakes. Most of the earthquakes occur on the big island of Hawaii, which is composed of the youngest, most recently built volcanoes. The geologic record shows that parts of the older islands have undergone major collapses in the last few million years, with sections of the islands sliding out to the seafloor in landslides floored on shallow normal faults.
Another example is the Basin and Range region of the western United States, including Nevada and eastern Utah, where the crust is subjected to tension. Earthquakes occur there on normal faults, far inland from the plate boundaries on the West Coast. The tension in the crust of the Basin and Range province may be partly due to a mid-ocean ridge system that subducted beneath California and is now located beneath the Basin and Range, causing tension in the lithosphere.
The region around Yellowstone National Park also undergoes occasional major earthquakes on normal faults. Earthquakes in that area may be due to the Yellowstone hot spot causing differential thermal expansion of the lithosphere in a broad zone round the hot spot center.
Several East Coast cities, including Boston, New York, and Charleston in South Carolina, have experienced damaging earthquakes in the last two centuries. The faults beneath these cities may date back to the rifting of Pangea and the opening up of the Atlantic Ocean beginning around 200 million years ago.
In the area of the town of New Madrid, along the Mississippi River in southeastern Missouri and western Tennessee, great earthquakes occurred in 1811-1812. Minor to moderate earthquakes continue to occur there, keeping active the possibility of damaging earthquakes occurring there again in the future. The fault system beneath that area may date from times of continental collision and continental rifting in the distant geologic past, and recent stress in the crust around New Madrid may be from the massive build-up of sediment in the Mississippi River delta region, which spreads out to the south of that area.
The connections between earthquakes and volcanoes are not always obvious. However, when magma is moving up beneath a volcano, and when a volcano is erupting, it produces earthquakes. Volcanic earthquakes are distinct from the more common type of earthquakes that occur by elastic rebound along faults.
Seismologists can use the patterns and signals of earthquakes coming from beneath volcanoes to predict that the volcano is about to erupt, and can use seismic waves to see that a volcano is undergoing an eruption even if the volcano is at a remote location, hidden in darkness, or hidden in storm clouds.
Volcanic vents, and volcanoes in general, are commonly located along faults, or at the intersection of several faults. Major faults that already exist in the crust may be natural paths to channel rising magma. However, on major volcanic edifices, shallower faults are a product of the development of the volcano. There are feedback effects between the upward pressure of magma buoyancy in the crust, the growth of faults in volcanic zones, and the venting of volcanoes, which is not yet completely understood.
As was noted at the beginning of this section, not quite all earthquakes are due to the slippage of solid blocks of rock along faults. When a volcano undergoes a powerful pyroclastic eruption - in other words, when a volcano explodes - it causes the earth to shake. Earthquakes caused by an explosive volcanic eruptions produce a different seismic signal than earthquakes caused by slippage along faults.
Another example of earthquakes that are caused at least in part by magma movement, rather than by slippage of entirely solid rock along faults, is earthquakes set off by the movement of magma upward beneath a volcano, or up to higher levels in the crust whether or not there is a volcano on top. Such upward movement of magma within the crust is sometimes called magma injection. Seismologists are still researching the interactions between movement of magma in the crust, and related slippage along faults that may be caused by the pressure and movement of the magma.
Earthquakes can be hazardous to humans and property in a variety of ways. Earthquake hazards arise from a combination of factors such as the size of the earthquake, distance to the epicenter, the underlying material and geologic structures, and building construction.
Ground shaking is caused by seismic waves During a significant earthquake, a particular location, and any building at that location, will be shaken by body waves (P- and S-) and surface waves (Rayleigh and Love. Each type of wave will have a different frequency (different number of waves passing by each second), and can shear or move a building in a different way, sometimes simultaneously. Adjacent neighborhoods or towns may experience very different intensities from the same earthquake, based on how far they are from the epicenter and what sort of rocks or sediments are in the ground beneath each area. Places underlain by thick deposits of unconsolidated sediments will experience a higher amplitude of shaking, at the same distance from the same earthquake, than places underlain by solid bedrock all the way to the surface. If unconsolidated sediments are fine-grained and wet, they may undergo liquefaction, increasing the damage to buildings and infrastructure and therefore increasing the intensity of the earthquake there. If unconsolidated sediments are overlain by a layer of artificial fill, the area is likely to experience more intense shaking and undergo more ground settling, and liquefaction if wet, than places that have not had a layer of artificial fill added.
Mexico City, one of the most populated cities in the world, is in a basin in the mountains of Mexico. Much of the city is built on artificial fill on top of fine-grained sediments from an extensive lake and wetlands that were drained and filled in as the city grew. As a result of how seismic waves are amplified in soft sediments, the shaking of the ground in Mexico City during an earthquake is greater than it is in areas outside the basin, which have bedrock close to the surface. During the 1985 earthquake, which originated offshore of Acapulco on the Pacific coast, 300 km (200 miles) away, many buildings in Mexico City collapsed and more than 20,000 people died.
The Marina District in northern San Francisco is built on artificial fill on wet bayshore sediments. Rubble and debris from buildings that collapsed or burned in that district as a result of the 1906 earthquake was used as artificial fill beneath structures built during the reconstruction of that area. In 1989, the Loma Prieta earthquake caused collapse of several buildings in the Marina District, and several people died there. Another example of a structure on wet ground covered by artificial fill that collapsed during the Loma Prieta earthquake is the Cypress viaduct in Oakland, CA. The USGS photo below shows the failed support columns. Click on thumbnail for a larger version that will open in a new window.
During large earthquakes, the ground may permanently shift to a new position up, down, or sideways (up to 10 or more m, 30 or 40 ft, in extreme cases). This change in the location of the ground, which also tilts the ground, may cause disruption of roads and utilities and, in coastal cities, submergence or emergence of harbor facilities. Even ground shifts of less than a meter (a foot or two) can cause serious disruption to infrastructure.
During most earthquakes, some rupturing of the earth's surface takes place along the fault trace. This produces a fault scarp, which may have up to several m (up to 10 ft or more) of vertical displacement. This can disrupt roads and utilities, and any buildings on a fault that ruptures may undergo extensive damage.
Rupture of the earth during an earthquake may also occur on secondary faults. Earthquake-induced rupture of the earth's surface may also take place in weak zones of surface sediment that fracture and spread. If enough spreading of a ruptured surface layer takes place, it can be classified as a landslide. The NOAA photo below shows substantial damage in the Turnagain-By-The-Sea subdivision caused by the 1964 Great Alaskan Earthquake. Click on thumbnail for a larger version that will open in a new window.
On steep slopes and in mountainous areas, large earthquakes can set off many landslides, rockfalls, or avalanches. These can damage buildings, towns, or roads in the path of the landslides.
If fine- or medium-grained, unconsolidated sediments are saturated with groundwater, the shaking that occurs during an earthquake may cause the sediment grains to lose contact with each other and become suspended in the water, temporarily turning what was solid ground into liquid ground. Building and other structures may sink, tilt, or slide a short distant in liquefied ground, causing serious damage.
Fires are a secondary rather than a primary effect of earthquakes. Broken electrical wires and natural gas pipes commonly set off fires during earthquakes. To compound the problem, the water supply may also be disrupted by earthquake damage, making it impossible to put the fire out with water from fire hydrants. The fire that broke out as a result of the great San Francisco earthquake of 1906 burned much of the city to the ground, causing more extensive damage to buildings than the shaking of the ground did during the earthquake.
A tsunami is a set of waves in the ocean (or a large lake) with an extremely long wavelength, typically over 100 km long. Tsunamis move at many 100s of km per hour in deep water. Tsunamis can be set off by violent volcanic eruptions originating just below sea level, by giant landslides that either occur underwater or tumble into the sea from coastal mountains, by large meteorite impacts, and, most commonly, by earthquakes that greatly shake the ocean floor, which commonly happens at subduction zones.
The amplitude, or crest height, of an individual tsunami wave may be only about 1 m (roughly 3 feet high) in the open ocean. It is common for a tsunami to pass ships at sea without being noticed. However, as the wave approaches shore where the bottom grows shallower, the crest builds up to a height of up to several tens of meters (over 30 feet in some cases). The wave crest may wash ashore for several minutes before subsiding. Even a tsunami wave of no higher than 3 m (10 feet) coming ashore can cause extensive damage in harbors and to shores, as the long-wavelength wave keeps pouring in for several minutes.
Tsunamis consist of more than one wave, so a second wave crest may climb ashore several minutes later. Some tsunamis lead with the wave trough, so the first thing noticed as that type of tsunami approaches a shore is a dramatic drawback, or retreat, of the sea, like the tide suddenly going out. Such a drawback will inevitably be followed by a rising tsunami wave crest.
The tsunamis from the Sumatra earthquake of 2004 killed over 100,000 people in coastal areas of the Indian Ocean, some on shores several thousand km (several thousand miles) away from the epicenter of the earthquake. The great southeast Alaska earthquake of 1964 generated a tsunami that killed 16 people on the coast of Oregon and northern California, over 1,000 km (600 miles) away. Tsunamis have been known to cross the entire Pacific Ocean and cause fatalities a third of the way around the world. Hawaii, in the middle of the Pacific Ocean, has been damaged by tsunamis, originating from subduction earthquakes on the Pacific Rim, several times in the last few centuries. Japan, with its complex set of subduction zones and its eastern shoreline open to the Pacific Ocean, has experienced over 100 tsunamis in its recorded history, most recently as a result of a great subduction earthquake off the northern island in 2011 which resulted in several coastal towns being destroyed and thousands of people dying. The last great tsunami generated by the Cascadia subduction zone along the coast of the Pacific Northwest flooded coastal areas of Washington, Oregon northern California, and Vancouver Island in Canada. It occurred in 1701, long before the region was as populated as it is now. Subduction continues in the Cascadia subduction zone and more tsunamis can be expected to be generated by earthquakes there in the future.
Much can be done to reduce the risk of fatalities during earthquakes, and to reduce the damage to buildings and infrastructure; in other words, to mitigate the effects of earthquakes.
In many cases, it is collapsed buildings that cause the most harm during an earthquake. Buildings should be constructed in ways that make them unlikely to collapse during an earthquake. The strategies that engineers have developed include having sufficient flexibility in the structure to absorb shaking during an earthquake. Bricks, mortar, and concrete are rigid and brittle. However, bricks and mortar, and concrete, can be reinforced with steel to make them better able to survive an earthquake. Wood and steel are more flexible than bricks, mortar, and concrete, and lend themselves to the type of building that, properly designed and built according to code, is likely to survive an earthquake without collapsing.
The way a building is attached to its foundation, and how the foundation is anchored in the earth, are important considerations in earthquake design. Many houses built in the the early and mid-1900s in California were not attached to their foundations, based on the assumption that the weight of a house would keep it on its foundation. It turned out to be a bad assumption. Earthquakes caused houses to slide off their foundations. Many home-owners in the state have taken steps to make sure that their houses are now attached to their foundations; if a home-owner buys earthquake insurance, it is usually required by the insurance company that they do so. A large building or skyscraper built in an earthquake-prone area will normally have a great deal of flexibility built into it, including some sort of elastic strain absorption mechanism focused on points where the building attaches to its foundation.
Infrastructure - roads, bridges, utilities - can be built with margins of safety for the event of an earthquake. This includes gas pipelines designed to slide back and forth on their supports and having built-in shutoff valves that may be activated by automatic sensors, electric lines and grids with similar flexibility and shut-off capabilities, and roads, overpasses, and bridges built to withstand shaking during an earthquake.
Development and enforcement of building codes aimed at reducing risk from earthquakes often requires resources that are not available in impoverished regions. This leads to a higher likelihood that buildings will collapse from the same size earthquake in some areas of the world than in other areas.
Seismicity is the study of how often earthquakes occur in a particular area, which types of earthquakes occur there, and why.
In the United States, the areas that most frequently experience earthquakes are the coast of California, Oregon, and Washington, the southern coast and Aleutian Islands of Alaska, Hawaii, and the mountain west from the Rocky Mountains to the Pacific coast. The central and eastern United States rarely experience significant earthquakes.
Earthquake epicenters compiled on a map show that, on a global basis, most earthquakes occur around the rim of the Pacific Ocean, in the mountains of southern Asia from China to the Middle East, and in the Mediterranean Sea area. Earthquake epicenters also trace the mid-ocean ridges across the floors of the oceans.
Because nearly all earthquakes occur on faults, determining seismic risks on a finer scale largely consists of identifying, mapping, and studying active faults in each state or region. However, many active faults are hidden, either because any scarps they formed at the surface have been eroded or covered by sediments, soil, and vegetation, or because they are blind faults. A hidden fault is often not identified and located until one or more significant earthquakes has occurred on it and the seismic waves have been studied to determine its location and type of fault motion.
Information used to determine the seismicity of an area includes:
- frequency of earthquakes in the past, as deduced from:
- historic records
- geologic studies that examine evidence of the prehistoric earthquake record
- location of known active faults
- seismologic data collected on recent earthquakes that have occurred in the area
- tectonic setting of the area in terms of proximity to plate boundaries, and information about the plate boundary if one is nearby
- stress and strain being experienced by the crust in that area based on measurements from GPS equipment and from stress and strain measurements conducted in boreholes
- underground geologic layers and structures in that area based on cross-sections from geologic mapping, data from drilling, and remote imaging of deeper layers of the crust and mantle
Based on this information, the seismic risk of a particular area can be quantified statistically. For example, the odds of a major earthquake happening in the next century, or in the next 10 years, can be estimated for a specific seismic zone.
However, no scientific method has yet been developed that can predict precisely when the next earthquake in a specific region will happen, where it will happen, or what its magnitude will be. Scientists have looked into using such possible pre-earthquake indicators as ground tilting, changes in well water levels, changes in radon gas in groundwater near fault zones, changes in electrical conductivity in the earth around faults, changes or patterns in seismic activity that can be measured by seismometers even though it is not felt by humans, and strange animal behavior which, according to numerous, largely unconfirmed anecdotes, takes place before an earthquake. But, so far, none of these types of data have been found to lead to reliable earthquake predictions.
People have also looked into correlations between earthquakes and phases of the Moon, earthquakes and the time of day (such as dawn when the Sun is first shining on the ground), and so on. No connections have been found between earthquakes and these other types of phenomena.
As the study of seismicity stands now, we can identify which areas on earth will undergo major earthquakes in the coming decades and centuries, we can delineate which areas on earth are at risk for the most powerful types of earthquakes, and map the coastal areas that are most at risk of being inundated by a tsunami, but we cannot pinpoint in advance the date or location of the next major earthquake.
Created by Ralph L. Dawes, Ph.D. and Cheryl D. Dawes, including figures unless otherwise noted
Unless otherwise specified, this work by Washington State Colleges is licensed under a Creative Commons Attribution 3.0 United States License.