The 1995 Kobe, Japan earthquake caused more than $100,000,000,000 in damage, and took more than 5000 lives.
We have explored the evidence for large-scale motions of the Earth's surface involved in the process of sea floor creation, sea floor subduction, and continental motions. Over the very long time scale processes involved, the Earth's mantle can actually be viewed as a viscous fluid, flowing convectively in the attempt to get heat out of the interior. But, because the surface is relatively cold, this fluid-like behavior is modified, stiffening the fluid, and resulting in the brittle behavior of rocks. The realm of brittle behavior gives rise to rock fracture and frictional sliding, which is the source of earthquakes.
How can rock flow like a fluid on one hand, yet break brittly on the other? This is somewhat counterintuitive, given our day to day experience with rocky materials, but think back and you will recall times where you drove by a road cut and observed strongly distorted rock structures, folded and bent from their original horizontal layering without having fractured. There are two key factors that must be considered: The effects of time, and the effects of temperature. Time is important, for even under the action of a steady, but fairly weak force, such as gravity at the surface of the Earth, a rock can begin to flow like a fluid given long enough time. This is observed in ancient human structures, where pillars and columns of Greek temples are abnormally thickened in their lower extremes, and stained glass windows in old European cathedrals are very thin at the top. Yet, if you rap on the stone or glass with a hammer, the material will chip or shatter. In this case, rapid time scale forces produce a different behavior than long time scale forces.
In addition, the behavior of materials changes with temperature. As one heats up a rock it will behave more and more ductile for both short and long time forces. The composition of the rock also influences its time and temperature behavior, so not all rocks have identical properties. We call these mechanical properties Rheology, meaning the behavior of solid materials in response to applied forces. Because the temperature increases with depth into the Earth, the rheology varies with depth. This is accentuated by the chemical variations between crust and mantle. The region of Earth that is very high viscosity, or stiff and brittle, is the low temperature zone near the surface called the Lithosphere. This is stiff even for long-term forces, but can bend like a beam. In fact, oceanic lithosphere need not fracture to deform as it eventually sinks back into the mantle. The lithosphere under oceans varies from very thin (about 5-10 km) near ridges to about 70-100 km in the oldest regions (150-200 my) of oceanic plate. The upper portion of the lithosphere is the crust, which is about 6 km thick in oceanic regions, but most of the lithosphere is upper mantle rocks. Note the distinction here, the crust is different than the mantle because it differs in rock composition (oceanic crust is basalt, while the upper mantle under it is peridotite depleted in basaltic components). The lithosphere is not defined by chemistry, but by rheology, in that it is the stiff region that translates coherently as a quasi-rigid plate. Under a continent the stiff region may extend 150-300 km deep, making for very thick regions of plate.
Within the lithosphere, heat transports by conduction, which does not involve physical flow of the material as in convection. The temperature increases almost linearly with depth in a conducting region, as in the lithosphere, but near temperatures of about 1000 degrees the rheology of the rock begins to change very rapidly and it becomes more ductile. The material then flows readily, and does not translate laterally with the moving plate, but shears. This softer region, which accommodates the motions of the lithosphere is called the asthenosphere. It is not molten rock (although there may be small amounts of partial melting in some regions), but for both short and long term forces the material flows rather than breaks. So, if you hit it with a hammer, it would be gushy, not brittle.
While the lithosphere is pretty stiff for long-term forces, throughout much of the lower part of the lithosphere, the rock is too warm to be brittle for short term forces. This restricts brittle frictional sliding and earthquake faulting to the coldest, shallowest regions of the lithosphere. Typically this is only the upper 20 km of continental rocks and the oceanic crust. It is possible for some earthquakes to occur in the uppermost mantle, although these are fairly rare. Deeper earthquakes require the special conditions of downwelling lithosphere to provide sufficiently brittle material to rupture at large depths.
Given the rheological properties that make the lithosphere stiff and its shallowest regions brittle for short term forces, the motions of plates are accompanied by shallow breaking and sliding of rock, rather than plastic deformation. This means that earthquakes occur on the planet, as the result of deformations in the shallowest layer that are responding to deep-seated convective motions as the planet cools. The ultimate source of energy for driving earthquakes is thus the heat of the Earth, and the convection processes that result in surface motions are responsible for earthquake activity.
To begin to consider earthquakes, the intrinsically brittle response of the upper portions of the lithosphere, we must define the idea of faults, which allow movements in the brittle regime. Faults are surfaces within the crust across which there have been shearing motions of the rocks on either side of the break. These surfaces are two-dimensional, and you can think of them as the surface contact between two halves of a broken stone.
Not all faults have earthquakes, as it is possible for rock to slide steadily and continuously, offsetting the two rock masses on either side of a fault with no sudden jarring motion. An earthquake is produced by sudden shearing slip on a fault, where sudden implies a time scale of fractions of a second (for the smallest earthquakes) up to as much as two minutes for the greatest earthquakes. This is a very short time relative to the time scale of plate tectonic motions and mantle convection, so earthquake faulting is on a more familiar time scale for humans than most plate tectonic phenomena. Of course, that is what makes the events so much more catastrophic.
Sudden, shearing slip on a fault is generally understood as a frictional sliding instability. This means that we view earthquake faulting as a process associated with frictional resistance to either breaking of rock (creating a new fault) or resistance to sliding of a previously broken rock with a fault in it. In either case, we can study friction in the laboratory to begin to understand what controls earthquake behavior. In general, friction on a fault increases the harder the rock is squeezed together (i.e. with the normal stress on the fault). More deeply buried rocks are more squeezed, so there is a general increase in frictional resistance with depth in the brittle zone. But the instability of friction proves to be very complex, as the conditions under which frictional resistance is suddenly overcome by a gradually building shearing stress on the rock depends on many properties such as the temperature, rock composition, presence of fluids in the rock and the connectedness of those fluids (which controls whether the fluids drain away or build up pore water pressure). The previous history of sliding of a fault proves to be important too, as each slip event grinds up rock, producing a layer of rock powder or gouge in the fault zone.
Faults exist on all scalelengths, ranging from hand-samples up to rock outcrops in roadcuts up to great breaks in the Earth's crust such as the San Andreas fault. The earthquake behavior of this huge range of faults varies accordingly. If we look at a map of where there are historically 'active' faults, meaning faults that have slipped in the most recent geological time interval called the Holocene, we find that California is criss-crossed with faults of many scales. There is a complex array of faults along the western margin of the state, with the single most continuous fault being the San Andreas Fault, which stretches from Cape Mendocino all the way to the Imperial Valley. It runs through the San Francisco peninsula, from just offshore of the Golden Gate Bridge, down through San Bernadino. This fault is the principle plate boundary fault between the Pacific and North American plates, but many other faults accommodate some of the relative motion between the plates (about 5 cm/yr for the Pacific/North American relative motions).
Eastern California has large faults as well, primarily on the eastern side of the Sierra Nevada. Whereas the San Andreas involves horizontal shearing of the crust, with the western side moving northwestward relative to the eastern side, the faults in the Owens Valley involve largely vertical motion, as the Sierra block tilts upward and the valley drops downward. There are a great variety of faults in California, as in all areas of the crust. Some are inactive, awaiting new patterns of crustal deformation to reactivate them or not, while others are quite active but are buried under sediments and we may be unaware of their existence.
On the largest scale, the plate boundaries of the Earth are all faults that connect up in various ways to make a world circuiting connection of breaks, across which the relative motions of the steadily moving plates are taking place. Since we have a pretty clear idea of how fast plates are moving (10s of cm/yr is typical), and from magnetic stripes, fault geometries, and active measurements of motions using lasers and satellite methods we know the current directions of relative plate motions, we have good constraints on how fast various faults are accumulating deformation that will be released in earthquakes.
For example, we know that the Pacific plate is moving northward at about 5 cm/yr relative to North America. This motion, driven by convection appears to be relatively steady over geological time periods. The edges of the plates are grinding past each other and must keep up with the overall motions, however, for most faults this is not a continuous sliding process, but a process of sticking and slipping, as friction resists sliding and is overcome in episodic events. Each sudden slip works to catch the plate offset up, but it may take multiple slips of a given strand of fault to catch up with the total plate offset over a given interval of time.
The San Andreas fault does not have uniform sliding properties, as there are long stretches of the fault that do not have small earthquakes, but appear to suddenly break every 100-150 years in very large earthquakes that involve 5-10 m of slip (thereby locally catching up with 100-150 years of relative plate motions at a steady rate of 5 cm/yr. There are other stretches of the fault that have multitudes of small earthquakes on a daily basis, apparently never suffering large ruptures. In some cases the offsets of the many small events may add up to keep up with the overall plate motion, but in some cases some of the motion is not accounted for by earthquakes, but involves steady sliding or creep of the fault. Clearly, the most catastrophic events are those that involve sudden great sliding of faults, but this tends to occur rarely, on a time scale that makes for a very limited record of past events upon which to base any prediction of the future behavior.
As we proceed to consider earthquake phenomena, we will need a few descriptive terms for describing faulting processes. These include:
If the faulting causes the upper block to move downward relative to the lower block (hanging wall downward relative to footwall), the fault is a normal fault. This is the type of fault usually found in regions of extension, where the crust is being pulled apart. This includes mid-ocean ridges where sea floor spreading is taking place as well as continental rifts, like in Eastern Africa, where the crust is breaking apart. The Basin and Range region of Eastern California, Nevada, Utah, and Idaho is also a region of normal faults due to regional extension and opening of the region. The faults on the eastern flank of the Sierra Nevada are normal faults.
If the hanging wall moves upward relative he footwall, you have reverse faulting, and if the dip (angle from the horizontal) of the fault is less than 30 degrees we call it thrust faulting. Reverse and thrust faulting occurs in regions of compression, where the surface is converging. This is common in subduction zones and in places where continents are colliding. The biggest thrust faults are those on the contact between underthrusting oceanic lithosphere and the overriding plate. The largest earthquakes tend to be thrust faulting events in subduction zones, and may be as large as magnitude 9.5.
The style of faulting reflects the regional tectonic process that is deforming the brittle crust, so it is very useful to be able to determine the faulting geometry whenever we can. In fact, we are able to do this without seeing the fault, as will be discussed later.
The 1906 San Francisco earthquake ruptured the northern 400 km of the San Andreas fault, with offsets of about 3-5 m along much of the length of the fault. This event caused a great destruction in San Francisco, in large part due to a conflagration that burned downtown. It was also an important event scientifically, as it involved the rupture of a vertically dipping right lateral strike slip fault, which offset the west side toward the northwest relative to the east side. This sense of displacement was readily observable in the fractured ground by offsets of fences, riverbeds, and other fault crossing features (including a railroad tunnel through the Santa Cruz mountains).
The clear observations of faulting and the intensive ensuing investigation of the earthquake prompted the articulation of the Elastic Rebound Theory, a conceptual model for how earthquake motions occur. The basic idea is that regional crustal movements, induced for example by plate tectonics, are acting over a region with a fault in it. If the fault were frictionless, the blocks of rock on either side would simply slide along steadily. Because of friction, the fault usually does not do this, and instead there is resistance to instantaneous motion. While rocks are brittle, they do have the ability to deform (strain) a tiny amount, as is true of any solid material. The rock on either side of the fault, in the so-called fault zone, strains to accommodate the regional motions, up to the point where the frictional resistance is exceeded. When this occurs, the rock slides on the fault surface, releasing the elastically accumulated strain energy in the fault zone, with the sudden fault slip catching up with the large scale offsets. The release of strain energy in the fault zone is fairly localized, controlled by the deformation properties of rock and the fact that fault zones are intrinsically weak regions in the crust because they are fractured.
Most of the energy released from the volume of strained rock around the fault goes into heat, as the sliding of rock heats up the fault surface. Some of the energy is released as seismic waves, which spread out through the rock, shaking the ground. It is the latter energy which causes most earthquake damage, as the waves expand outward and vibrate human structures. Fairly little damage occurs as the direct result of abrupt offsets in the fault zone, although there is a tendency for humans to build in fault zone regions (either they have carved convenient valleys and riverways which attract people, or there are exposed cliffs adjacent to fault scarps which have nice views).
Understanding the seismic waves is key to understanding how earthquakes cause catastrophes. There are 2 fundamental seismic waves. P (Primary) waves and S (Secondary) waves. P waves travel faster than S waves and are directly analogous to sound waves in a solid material. We are familiar with sound properties, which involve the excitation of air molecule vibrations by a source (such as a vibrating vocal cord), and the outward propagation of that sound as adjacent molecules are vibrated and then more distant ones vibrate etc., until finally some vibrate a sensor such as your eardrum and your brain translates the electrical impulses from the eardrum into an awareness of the disturbance. In sound waves the particles oscillate back and forth in the direction that the sound propagates (outward in all directions from the source on a spherical wavefront). After the sound passes the air particles return to where they were before the sound wave passed through the medium. Similar oscillations occur in a P wave in rock, in that the particles oscillate back and forth in the same direction as the wave is propagating, returning to their original position due to the restoring forces of the surrounding rock. When we hear a noise through a wall, we are simply hearing a sound that propagated through air, turned to a P wave as it went through the solid wall, and again became a sound wave that we eventually hear.
S waves involve shearing motions perpendicular to the direction in which the wave disturbance is propagating. In a rock, the adjacent material has a restoring force that causes the shearing particles to return elastically to their original position. If you try to shear a fluid, there is no effective restoring force, so S waves cannot propagate in water or air.
It is important to remember that waves spread in all directions away from the source, effectively on spherical wavefronts surrounding the source in the rock. So, at an instant of time after an earthquake fault slides, the source will be surrounded by an outward propagating P wave and an outward propagating S wave. The velocity of the P wave, or the velocity of sound in a rock, is faster so the P wavefront spreads through the rock faster. The P and S velocities are both controlled by material properties of the rock, in particular the P velocity is controlled by the bulk modulus or compressibility of the rock, its rigidity (resistance to shear), and its density. The shear velocity is controlled by rigidity and density.
The outward propagating P and S waves spread through the Earth, with the amplitude of the wave decreasing as the wave travels further. This is because the energy is spread over a larger and larger surface as a function of increasing time. Eventually, the wave is too small to detect. The larger the initial input of energy (i.e. the larger the earthquake), the more distant the perceptible shaking will be. With sensitive instruments we can detect the tiny motions of small earthquakes (say, magnitude 4.5) everywhere on the Earth, even though these events don't usually cause damage even right at the fault region. A large earthquake, such as a magnitude 7 earthquake, may cause destructive shaking for tens of miles, and the waves can be detecting circling the planet again and again before they die down to imperceptible levels. A great, magnitude 9 earthquake can set the Earth to ringing like a bell for days.
We study the seismic waves from earthquakes because they tell us about the source (the earthquake faulting process) and they tell us about the Earth. Major results from the study of seismic waves include:
We will explore all of these.
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