Magma

Magma is molten or partially molten rock beneath the Earth’s surface. When magma erupts onto the surface, it is called lava. Magma typically consists of (1) a liquid portion (often referred to as the melt); (2) a solid portion made of minerals that crystallized directly from the melt; (3) solid rocks incorporated into the magma from along the conduit or reservoir, called xenoliths or inclusions; and (4) dissolved gases.

Magma (Plurals: magmas and magmata) is molten rock that sometimes forms beneath the surface of the Earth (or any other terrestrial planet) that often collects in a magma chamber. Magma may contain suspended crystals and gas bubbles. By definition, all igneous rock is formed from magma.

Hawaiian lava flow (lava is the extrusive equivalent of magma *Pahoehoe)

Magma is a complex high-temperature fluid substance. Temperatures of most magmas are in the range 700°C to 1300°C, but very rare carbonatite melts may be as cool as 600°C, and komatiite melts may have been as hot at 1600°C. Most are silicate solutions.

It is capable of intrusion into adjacent rocks or of extrusion onto the surface as lava or ejected explosively as tephra to form pyroclastic rock.

Environments of magma formation and compositions are commonly correlated. Environments include subduction zones, continental rift zones, mid-oceanic ridges, and hotspots, some of which are interpreted as mantle plumes. Environments are discussed in the entry on igneous rock. Magma compositions may evolve after formation by fractional crystallization, contamination, and magma mixing.

Contrary to some impressions,the bulk of the Earth’s crust and mantle is not molten. Rather, the bulk of the Earth takes the form of a rheid, a form of solid that can move or deform under pressure. Magma, as liquid, preferentally forms in high temperature, low pressure environments within several kilometers of the Earth’s surface.

Melting of solid rock

Melting of solid rock to form magma is controlled by three physical parameters: its temperature, pressure, and composition. Mechanisms are discussed in the entry for igneous rock.

Temperature

At any given pressure and for any given composition of rock, a rise in temperature past the solidus will cause melting. Within the solid earth, the temperature of a rock is controlled by the geothermal gradient and the radioactive decay within the rock. The geothermal gradient averages about 25°C/km with a wide range from a low of 5-10°C/km within oceanic trenches and subduction zones to 30-80°C/km under mid-ocean ridges and volcanic arc environments.

Pressure

Melting can also occur due to a reduction in pressure by a process known as decompression melting.^[1]

Composition

It is usually very difficult to change the bulk composition of a large mass of rock, so composition is the basic control on whether a rock will melt at any given temperature and pressure. The composition of a rock may also be considered to include volatile phases such as water and carbon dioxide.

The presence of volatile phases in a rock under pressure can stabilize a melt fraction. The presence of even 0.8% water may reduce the temperature of melting by as much as 100°C. Conversely, the loss of water and volatiles from a magma may cause it to essentially freeze or solidify.

 Partial melting

When rocks melt they do so incrementally and gradually; most rocks are made of several minerals, all of which have different melting points, and the phase diagrams that control melting commonly are complex. As a rock melts, its volume changes. When enough rock is melted, the small globules of melt (generally occurring in between mineral grains) link up and soften the rock. Under pressure within the earth, as little as a fraction of a percent partial melting may be sufficient to cause melt to be squeezed from its source.

Melts can stay in place long enough to melt to 20% or even 35%, but rocks are rarely melted in excess of 50%, because eventually the melted rock mass becomes a crystal and melt mush that can then ascend en masse as a diapir, which may then cause further decompression melting.

 Primary melts

When a rock melts, the liquid is known as a primary melt. Primary melts have not undergone any differentiation and represent the starting composition of a magma. In nature it is rare to find primary melts. The leucosomes of migmatites are examples of primary melts. Primary melts derived from the mantle are especially important, and are known as primitive melts or primitive magmas. By finding the primitive magma composition of a magma series it is possible to model the composition of the mantle from which a melt was formed, which is important in understanding evolution of the mantle.

 Parental melts

Where it is impossible to find the primitive or primary magma composition, it is often useful to attempt to identify a parental melt. A parental melt is a magma composition from which the observed range of magma chemistries has been derived by the processes of igneous differentiation. It need not be a primitive melt.

For instance, a series of basalt flows are assumed to be related to one another. A composition from which they could reasonably be produced by fractional crystallization is termed a parental melt. Fractional crystallization models would be produced to test the hypothesis that they share a common parental melt.

 Geochemical implications of partial melting

The degree of partial melting is critical for determining what type of magma is produced. The degree of partial melting required to form a melt can be estimated by considering the relative enrichment of incompatible elements versus compatible elements. Incompatible elements commonly include potassium, barium, caesium, rubidium.

Rock types produced by small degrees of partial melting in the Earth’s mantle are typically alkaline (Ca, Na), potassic (K) and/or peralkaline (high aluminium to silica ratio). Typically, primitive melts of this composition form lamprophyre, lamproite, kimberlite and sometimes nepheline-bearing mafic rocks such as alkali basalts and essexite gabbros or even carbonatite.

Pegmatite may be produced by low degrees of partial melting of the crust. Some granite-composition magmas are eutectic (or cotectic) melts, and they may be produced by low to high degrees of partial melting of the crust, as well as by fractional crystallization. At high degrees of partial melting of the crust, granitoids such as tonalite, granodiorite and monzonite can be produced, but other mechanisms are typically important in producing them.

At high degrees of partial melting of the mantle, komatiite and picrite are produced.

Composition and melt structure and properties

Silicate melts are composed mainly of silicon, oxygen, aluminium, alkalis (sodium, potassium, calcium), magnesium and iron. Silicon atoms are in tetrahedral coordination with oxygen, as in almost all silicate minerals, but in melts atomic order is preserved only over short distances. The physical behaviours of melts depend upon their atomic structures as well as upon temperature and pressure and composition.

Viscosity is a key melt property in understanding the behaviour of magmas. More silica-rich melts are typically more polymerized, with more linkage of silica tetrahedra, and so are more viscous. Dissolution of water drastically reduces melt viscosity. Higher-temperature melts are less viscous.

Generally speaking, more mafic magmas, such as those that form basalt, are hotter and less viscous than more silica-rich magmas, such as those that form rhyolite. Low viscosity leads to gentler, less explosive eruptions.

Main article: Igneous differentiation

Characteristics of several different magma types are as follows:

Ultramafic (picritic)

SiO₂ < 45%

Fe-Mg >8% up to 32%MgO

Temperature: up to 1500°C

Viscosity: Very Low

Eruptive behavior: gentle or very explosive (kimberilites)

Distribution: divergent plate boundaries, hot spots, convergent plate boundaries; komatiite and other ultramafic lavas are mostly Archean and were formed from a higher geothermal gradient and are unknown in the present

Mafic (basaltic)

SiO₂ < 50%

FeO and MgO typically < 10 wt%

Temperature: up to ~1300°C

Viscosity: Low

Eruptive behavior: gentle

Distribution: divergent plate boundaries, hot spots, convergent plate boundaries

Intermediate (andesitic)

SiO₂ ~ 60%

Fe-Mg: ~ 3%

Temperature: ~1000°C

Viscosity: Intermediate

Eruptive behavior: explosive

Distribution: convergent plate boundaries

Felsic (rhyolitic)

SiO₂ >70%

Fe-Mg: ~ 2%

Temp: < 900°C

Viscosity: High

Eruptive behavior: explosive

Distribution: hot spots in continental crust (Yellowstone National Park), continental rifts, island arcs

See also

• Igneous differentiation
• Cumulate rocks
• Fractional crystallization
• Igneous rocks
• Lava
• Intrusion
• Ultramafic to mafic layered intrusions

source:http://en.wikipedia.org/wiki/Magma

VOLCANO TYPES

GENERIC FEATURES

A volcanic vent is an opening exposed on the earth’s surface where volcanic material is emitted. All volcanoes contain a central vent underlying the summit crater of the volcano. The volcano’s cone-shaped structure, or edifice, is built by the more-or-less symmetrical accumulation of lava and/or pyroclastic material around this central vent system. The central vent is connected at depth to a magma chamber, which is the main storage area for the eruptive material. Because volcano flanks are inherently unstable, they often contain fractures that descend downward toward the central vent, or toward a shallow-level magma chamber. Such fractures may occasionally tap the magma source and act as conduits for flank eruptions along the sides of the volcanic edifice. These eruptions can generate cone-shaped accumulations of volcanic material, called parasitic cones. Fractures can also act as conduits for escaping volcanic gases, which are released at the surface through vent openings called fumaroles.

Summit Crater

Parasitic Cones

Fumarole

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MAIN VOLCANO TYPES

Although every volcano has a unique eruptive history, most can be grouped into three main types based largely on their eruptive patterns and their general forms. The form and composition of the three main volcano types are summarized here:

VOLCANO TYPE	VOLCANO SHAPE	COMPOSITION	ERUPTION TYPE
SCORIA CONE	Straight sides with steep slopes; large summit crater	Basalt tephra; occasionally andesitic	Strombolian
SHIELD VOLCANO	Very gentle slopes; convex upward	Basalt lava flows	Hawaiian
STRATOVOLCANO	Gentle lower slopes, but steep upper slopes; concave upward; small summit crater	Highly variable; alternating basaltic to rhyolitic lavas and tephra with an overall andesite composition	Plinian

SUBORDINATE VOLCANO TYPES — Lava and tephra can erupt from vents other than these three main volcano types. A fissure eruption, for example, can generate huge volumes of basalt lava; however, this type of eruption is not associated with the construction of a volcanic edifice around a single central vent system. Although point-source eruptions can generate such features as spatter cones and hornitos, these volcanic edifices are typically small, localized, and/or associated with rootless eruptions (i.e., eruptions above the surface of an active lavaflow, unconnected to an overlying magma chamber) . Vent types related to hydrovolcanic processes generate unique volcanic structures, discussed separately under hydrovolcanic eruptions.

For a description of each of the main volcano types, see:

• SCORIA CONES
• SHIELD VOLCANOES
• STRATOVOLCANOES

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WHEN IS A VOLCANO CONSIDERED ACTIVE, DORMANT, OR EXTINCT?

Classifying a volcano as active, dormant, or extinct is a subjective and inexact exercise. A volcano is generally considered active if it has erupted in historic time. This definition, however, is rather ambiguous, because recorded history varies from thousands of years in Europe and the Middle East, to only a few hundred years in other regions of the world, like the Pacific Northwest of the United States. Scientists generally consider a volcano active if it is currently erupting, or exhibiting unrest through earthquakes, uplift, and/or new gas emissions. The Smithsonian Institution’s catalog of active volcanoes, recognizes 539 volcanoes with historic eruptions. In addition, there are 529 volcanoes that have not erupted in historic times, but which exhibit clear evidence of eruption in the past 10,000 years. These latter volcanoes are probably best considered “dormant,” since they have the potential to erupt again.

Whether or not inactive volcanoes are considered truly extinct, or just dormant, depends partly on the average repose interval between eruptions. As noted in eruptive variability, explosive eruptions like those at Toba and Yellowstone have repose intervals of hundreds of thousands of years, whereas non-explosive eruptions have very short repose intervals. Thus, the Yellowstone region, which has not experienced an eruption for 70,000 years, can not be considered extinct. In fact, many scientists consider Yellowstone to be active because of high uplift rates, frequent earthquakes, and a very active geothermal system. Many inactive scoria cones, on the other hand, may be viewed as extinct shortly after they erupt, because such volcanoes are typically monogenetic and only erupt once.

 source:http://www.geology.sdsu.edu

The Modified Mercalli Scale

In seismology a scale of seismic intensity is a way of measuring or rating the effects of an earthquake at different sites. The Modified Mercalli Intensity Scale is commonly used in the United States by seismologists seeking information on the severity of earthquake effects. Intensity ratings are expressed as Roman numerals between I at the low end and XII at the high end. The Intensity Scale differs from the Richter Magnitude Scale in that the effects of any one earthquake vary greatly from place to place, so there may be many Intensity values (e.g.: IV, VII) measured from one earthquake. Each earthquake, on the other hand, should have just one Magnitude, although the several methods of estimating it will yield slightly different values (e.g.: 6.1, 6.3).

Ratings of earthquake effects are based on the following relatively subjective scale of descriptions:

Modified Mercalli Intensity Scale

from FEMA

I. People do not feel any Earth movement.

II. A few people might notice movement if they are at rest and/or on the upper floors of tall buildings.

III. Many people indoors feel movement. Hanging objects swing back and forth. People outdoors might not realize that an earthquake is occurring.

IV. Most people indoors feel movement. Hanging objects swing. Dishes, windows, and doors rattle. The earthquake feels like a heavy truck hitting the walls. A few people outdoors may feel movement. Parked cars rock.

V. Almost everyone feels movement. Sleeping people are awakened. Doors swing open or close. Dishes are broken. Pictures on the wall move. Small objects move or are turned over. Trees might shake. Liquids might spill out of open containers.

VI. Everyone feels movement. People have trouble walking. Objects fall from shelves. Pictures fall off walls. Furniture moves. Plaster in walls might crack. Trees and bushes shake. Damage is slight in poorly built buildings. No structural damage.

VII. People have difficulty standing. Drivers feel their cars shaking. Some furniture breaks. Loose bricks fall from buildings. Damage is slight to moderate in well-built buildings; considerable in poorly built buildings.

VIII. Drivers have trouble steering. Houses that are not bolted down might shift on their foundations. Tall structures such as towers and chimneys might twist and fall. Well-built buildings suffer slight damage. Poorly built structures suffer severe damage. Tree branches break. Hillsides might crack if the ground is wet. Water levels in wells might change.

IX. Well-built buildings suffer considerable damage. Houses that are not bolted down move off their foundations. Some underground pipes are broken. The ground cracks. Reservoirs suffer serious damage.

X. Most buildings and their foundations are destroyed. Some bridges are destroyed. Dams are seriously damaged. Large landslides occur. Water is thrown on the banks of canals, rivers, lakes. The ground cracks in large areas. Railroad tracks are bent slightly.

XI. Most buildings collapse. Some bridges are destroyed. Large cracks appear in the ground. Underground pipelines are destroyed. Railroad tracks are badly bent.

XII. Almost everything is destroyed. Objects are thrown into the air. The ground moves in waves or ripples. Large amounts of rock may move.

As you can see from the list above, rating the Intensity of an earthquake’s effects does not require any instrumental measurements. Thus seismologists can use newspaper accounts, diaries, and other historical records to make intensity ratings of past earthquakes, for which there are no instrumental recordings. Such research helps promote our understanding of the earthquake history of a region, and estimate future hazards.

This map plots the Mercalli Intensity ratings of localities near the Oct. 17, 1989 Loma Prieta (World Series) earthquake. It is called an isoseismal map, as one draws contour lines to enclose locations having higher intensities. Intensities typically increase close to an earthquake’s epicenter, allowing seismologists to interpret maps such as this for the general location of historical earthquakes.

Note the locations of unusually high intensities (up to IX) far north of the earthquake’s epicenter, near San Francisco Bay. During this earthquake, soft and water-saturated soils near the Bay amplified the effects of the shaking. The amplified shaking, together with soil liquefaction effects, caused some well-built structures to collapse and yielded the intensity IX rating at those locations.

It is also possible to estimate the Magnitude of an earthquake from the area of the map enclosed by isoseismal contours of certain intensities. Such estimates are, however, a subject of research and require verification.

source:http://www.seismo.unr.edu

Richter Scale

What is Richter Magnitude?

Short answer:

Seismologists use a Magnitude scale to express the seismic energy released by each earthquake. Here are the typical effects of earthquakes in various magnitude ranges:

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Earthquake Severity

Richter         Earthquake
Magnitudes      Effects

Less than 3.5   Generally not felt, but recorded.

3.5-5.4         Often felt, but rarely causes damage.

Under 6.0       At most slight damage to well-designed buildings.
		Can cause major damage to poorly constructed buildings
		over small regions.

6.1-6.9         Can be destructive in areas up to about 100 kilometers
		across where people live.

7.0-7.9         Major earthquake. Can cause serious damage over larger areas.

8 or greater    Great earthquake. Can cause serious damage in areas several
		hundred kilometers across.

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Although each earthquake has a unique Magnitude, its effects will vary greatly according to distance, ground conditions, construction standards, and other factors. Seismologists use a different Mercalli Intensity Scale to express the variable effects of an earthquake. Each earthquake has a unique amount of energy, but magnitude values given by different seismological observatories for an event may vary. Depending on the size, nature, and location of an earthquake, seismologists use several different methods to estimate magnitude. The uncertainty in an estimate of the magnitude is about plus or minus 0.3 units, and seismologists often revise magnitude estimates as they obtain and analyze additional data.

With permission from http://www.comics.com/comics/franknernest/index.html <!– Click here for a funny Frank and Earnest comic about seismologists, by Bob Thaves. –>

Long answer:

One of Dr. Charles F. Richter’s most valuable contributions was to recognize that the seismic waves radiated by all earthquakes can provide good estimates of their magnitudes. You can read about seismic waves by clicking here. He collected the recordings of seismic waves from a large number of earthquakes, and developed a calibrated system of measuring them for magnitude. Richter showed that, the larger the intrinsic energy of the earthquake, the larger the amplitude of ground motion at a given distance. He calibrated his scale of magnitudes using measured maximum amplitudes of shear waves on seismometers particularly sensitive to shear waves with periods of about one second. The records had to be obtained from a specific kind of instrument, called a Wood-Anderson seismograph. Although his work was originally calibrated only for these specific seismometers, and only for earthquakes in southern California, seismologists have developed scale factors to extend Richter’s magnitude scale to many other types of measurements on all types of seismometers, all over the world. In fact, magnitude estimates have been made for thousands of Moon-quakes and for two quakes on Mars.

The diagram below demonstrates how to use Richter’s original method to measure a seismogram for a magnitude estimate in Southern California:

The scales in the diagram above form a nomogram that allows you to do the mathematical computation quickly by eye. The equation for Richter Magnitude is:

M_L = log₁₀A(mm) + (Distance correction factor)
Here A is the amplitude, in millimeters, measured directly from the photographic paper record of the Wood-Anderson seismometer, a special type of instrument. The distance factor comes from a table that can be found in Richter’s (195 book Elementary Seismology. The equation behind this nomogram, used by Richter in Southern California, is:

Thus after you measure the wave amplitude you have to take its logarithm, and scale it according to the distance of the seismometer from the earthquake, estimated by the S-P time difference. The S-P time, in seconds, makes .

Click here to learn more about the mathematical logarithm.

Seismologists will try to get a separate magnitude estimate from every seismograph station that records the earthquake, and then average them. This accounts for the usual spread of around 0.2 magnitude units that you see reported from different seismological labs right after an earthquake. Each lab is averaging in different stations that they have access to. It may be several days before different organizations will come to a consensus on what was the best magnitude estimate.

Seismic Moment:

Seismologists have more recently developed a standard magnitude scale that is completely independent of the type of instrument. It is called the moment magnitude, and it comes from the seismic moment. To get an idea of the seismic moment, we go back to the elementary physics concept of torque. A torque is a force that changes the angular momentum of a system. It is defined as the force times the distance from the center of rotation. Earthquakes are caused by internal torques, from the interactions of different blocks of the earth on opposite sides of faults. After some rather complicated mathematics, it can be shown that the moment of an earthquake is simply expressed by:

The formula above, for the moment of an earthquake, is fundamental to seismologists’ understanding of how dangerous faults of a certain size can be.

Now, let’s imagine a chunk of rock on a lab bench, the rigidity, or resistance to shearing, of the rock is a pressure in the neighborhood of a few hundred billion dynes per square centimeter. (Scientific notation makes this easier to write.) The pressure acts over an area to produce a force, and you can see that the cm-squared units cancel. Now if we guess that the distance the two parts grind together before they fly apart is about a centimeter, then we can calculate the moment, in dyne-cm:

Again it is helpful to use scientific notation, since a dyne-cm is really a puny amount of moment.

Now let’s consider a second case, the Sept. 12, 1994 Double Spring Flat earthquake, which occurred about 25 km southeast of Gardnerville. The first thing we have to do, since we’re working in centimeters, is figure out how to convert the 15 kilometer length and 10 km depth of that fault to centimeters. We know that 100 thousand centimeters equal one kilometer, so we can write that equation and divide both sides by “km” to get a factor equal to one.

Of course we can multiply anything by one without changing it, so we use it to cancel the kilometer units and put in the right centimeter units:

Of course this result needs scientific notation even more desperately. We can see that this earthquake, the largest in Nevada in 28 years, had two times ten raised to the twelfth power, or 2 trillion, times as much moment as breaking the rock on the lab table.

There is a standard way to convert a seismic moment to a magnitude. The equation is:

Now let’s use this equation (meant for energies expressed in dyne-cm units) to estimate the magnitude of the tiny earthquake we can make on a lab table:

Negative magnitudes are allowed on Richter’s scale, although such earthquakes are certainly very small.

Next let’s take the energy we found for the Double Spring Flat earthquake and estimate its magnitude:

The magnitude 6.1 value we get is about equal to the magnitude reported by the UNR Seismological Lab, and by other observers.

Seismic Energy:

Both the magnitude and the seismic moment are related to the amount of energy that is radiated by an earthquake. Richter, working with Dr. Beno Gutenberg, early on developed a relationship between magnitude and energy. Their relationship is: logE_S = 11.8 + 1.5M
giving the energy E_S in ergs from the magnitude M. Note that E_S is not the total “intrinsic” energy of the earthquake, transferred from sources such as gravitational energy or to sinks such as heat energy. It is only the amount radiated from the earthquake as seismic waves, which ought to be a small fraction of the total energy transfered during the earthquake process.

More recently, Dr. Hiroo Kanamori came up with a relationship between seismic moment and seismic wave energy. It gives:

Energy = (Moment)/20,000
For this moment is in units of dyne-cm, and energy is in units of ergs. dyne-cm and ergs are unit equivalents, but have different physical meaning.

Let’s take a look at the seismic wave energy yielded by our two examples, in comparison to that of a number of earthquakes and other phenomena. For this we’ll use a larger unit of energy, the seismic energy yield of quantities of the explosive TNT (We assume one ounce of TNT exploded below ground yields 640 million ergs of seismic wave energy):

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Richter     TNT for Seismic    Example
Magnitude      Energy Yield    (approximate)

-1.5                6 ounces   Breaking a rock on a lab table
 1.0               30 pounds   Large Blast at a Construction Site
 1.5              320 pounds
 2.0                1 ton      Large Quarry or Mine Blast
 2.5              4.6 tons
 3.0               29 tons
 3.5               73 tons
 4.0            1,000 tons     Small Nuclear Weapon
 4.5            5,100 tons     Average Tornado (total energy)
 5.0           32,000 tons
 5.5           80,000 tons     Little Skull Mtn., NV Quake, 1992
 6.0        1 million tons     Double Spring Flat, NV Quake, 1994
 6.5        5 million tons     Northridge, CA Quake, 1994
 7.0       32 million tons     Hyogo-Ken Nanbu, Japan Quake, 1995; Largest Thermonuclear Weapon
 7.5      160 million tons     Landers, CA Quake, 1992
 8.0        1 billion tons     San Francisco, CA Quake, 1906
 8.5        5 billion tons     Anchorage, AK Quake, 1964
 9.0       32 billion tons     Chilean Quake, 1960
10.0       1 trillion tons     (San-Andreas type fault circling Earth)
12.0     160 trillion tons     (Fault Earth in half through center,
                               OR Earth's daily receipt of solar energy)

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160 trillion tons of dynamite is a frightening yield of energy. Consider, however, that the Earth receives that amount in sunlight every day.

Practical ways of estimating magnitude

Most seismologists prefer to use the seismic moment to estimate earthquake magnitudes. Finding an earthquake fault’s length, depth, and its slip can take several days, weeks, or even months after a big earthquake. Geologists’ mapping of the earthquake’s fault breaks, or seismologists’ plotting of the spatial distribution of aftershocks, can give these parameters after a substantial effort. But some large earthquakes, and most small earthquakes, show neither surface fault breaks nor enough aftershocks to estimate magnitudes the way we have above. However, seismologists have developed ways to estimate the seismic moment directly from seismograms using computer processing methods. The Centroid Moment Tensor Project at Harvard University has been routinely estimating moments of large earthquakes around the world by seismogram inversion since 1982.

Another measure of an earthquake

Seismologists use a separate method to estimate the effects of an earthquake, called its intensity. Intensity should not be confused with magnitude. Although each earthquake has a single magnitude value, its effects will vary from place to place, and there will be many different intensity estimates. You can read about the Mercalli Intensity Scale, one popular way to characterize earthquake effects.

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source: http://www.seismo.unr.edu

Seismic Waves

Seismic Deformation

When an earthquake fault ruptures, it causes two types of deformation: static; and dynamic. Static deformation is the permanent displacement of the ground due to the event. The earthquake cycle progresses from a fault that is not under stress, to a stressed fault as the plate tectonic motions driving the fault slowly proceed, to rupture during an earthquake and a newly-relaxed but deformed state.

Typically, someone will build a straight reference line such as a road, railroad, pole line, or fence line across the fault while it is in the pre-rupture stressed state. After the earthquake, the formerly stright line is distorted into a shape having increasing displacement near the fault, a process known as elastic rebound.

Seismic Waves

The second type of deformation, dynamic motions, are essentially sound waves radiated from the earthquake as it ruptures. While most of the plate-tectonic energy driving fault ruptures is taken up by static deformation, up to 10% may dissipate immediately in the form of seismic waves.

The mechanical properties of the rocks that seismic waves travel through quickly organize the waves into two types. Compressional waves, also known as primary or P waves, travel fastest, at speeds between 1.5 and 8 kilometers per second in the Earth’s crust. Shear waves, also known as secondary or S waves, travel more slowly, usually at 60% to 70% of the speed of P waves.

P waves shake the ground in the direction they are propagating, while S waves shake perpendicularly or transverse to the direction of propagation.

Although wave speeds vary by a factor of ten or more in the Earth, the ratio between the average speeds of a P wave and of its following S wave is quite constant. This fact enables seismologists to simply time the delay between the arrival of the P wave and the arrival of the S wave to get a quick and reasonably accurate estimate of the distance of the earthquake from the observation station. Just multiply the S-minus-P (S-P) time, in seconds, by the factor 8 km/s to get the approximate distance in kilometers.

The dynamic, transient seismic waves from any substantial earthquake will propagate all around and entirely through the Earth. Given a sensitive enough detector, it is possible to record the seismic waves from even minor events occurring anywhere in the world at any other location on the globe. Nuclear test-ban treaties in effect today rely on our ability to detect a nuclear explosion anywhere equivalent to an earthquake as small as Richter Magnitude 3.5.

Seismographs and Seismograms

Sensitive seismographs are the principal tool of scientists who study earthquakes. Thousands of seismograph stations are in operation throughout the world, and instruments have been transported to the Moon, Mars, and Venus. Fundamentally, a seismograph is a simple pendulum. When the ground shakes, the base and frame of the instrument move with it, but intertia keeps the pendulum bob in place. It will then appear to move, relative to the shaking ground. As it moves it records the pendulum displacements as they change with time, tracing out a record called a seismogram.
One seismograph station, having three different pendulums sensitive to the north-south, east-west, and vertical motions of the ground, will record seismograms that allow scientists to estimate the distance, direction, Richter Magnitude, and type of faulting of the earthquake. Seismologists use networks of seismograph stations to determine the location of an earthquake, and better estimate its other parameters. It is often revealing to examine seismograms recorded at a range of distances from an earthquake:

On this example it is obvious that seismic waves take more time to arrive at stations that are farther away. The average velocity of the wave is just the slope of the line connecting arrivals, or the change in distance divided by the change in time. Variations in such slopes reveal variations in the seismic velocities of rocks. Note the secondary S-wave arrivals that have larger amplitudes than the first P waves, and connect at a smaller slope. While the actual frequencies of seismic waves are below the range of human hearing, it is possible to speed up a recorded seismogram to hear it. You can click on this earthquake recording to hear a seismogram from the 1992 Landers earthquake in southern California, recorded near Mammoth Lakes in an active volcanic caldera by the USGS. The original record, 800 seconds long, has been speeded up 80 times so that you hear it all within 10 seconds.

75 kb u-law; 149 kb WAV; 75 kb Quicktime The clicks at the beginning of the recording are the sharp, high-frequency P waves, followed by the rushing sound of the drawn-out, lower-frequency S waves. This recording is also interesting because of the small, local earthquakes within the Mammoth caldera that sound like gunshots. The passage of the S wave from the magnitude 7.2 Landers event through the caldera actually triggered a sequence of small earthquakes there. The triggered earthquakes are similar to a burst of creaks and pops you hear from your house frame after a strong blast of wind. Landers triggered earthquakes up to magnitude 5.5 throughout eastern California and Nevada, and in calderas as far away as Yellowstone.

Listen to more earthquakes with:

• John Louie’s The Sound of Seismic (MP3s)
• Andy Michael and Dennis Ross’s Listening to Earthquakes (WAVs)

Locating Earthquakes

The pricipal use of seismograph networks is to locate earthquakes. Although it is possible to infer a general location for an event from the records of a single station, it is most accurate to use three or more stations. Locating the source of any earthquake is important, of course, in assessing the damage that the event may have caused, and in relating the earthquake to its geologic setting. Given a single seismic station, the seismogram records will yield a measurement of the S-P time, and thus the distance between the station and the event. Multiply the seconds of S-P time by 8 km/s for the kilometers of distance. Drawing a circle on a map around the station’s location, with a radius equal to the distance, shows all possible locations for the event. With the S-P time from a second station, the circle around that station will narrow the possible locations down to two points. It is only with a third station’s S-P time that you can draw a third circle that should identify which of the two previous possible points is the real one:
This example uses stations in Boston, Edinborough, and Manaus. With the distances shown, all three circles can intersect only at a single point on the Mid-Atlantic Ridge spreading cen

source:http://www.seismo.unr.edu

Plate Tectonics, the Cause of Earthquakes

The plates consist of an outer layer of the Earth, the lithosphere, which is cool enough to behave as a more or less rigid shell. Occasionally the hot asthenosphere of the Earth finds a weak place in the lithosphere to rise buoyantly as a plume, or hotspot. The satellite image below shows the volcanic islands of the Galapagos hotspot.

(from NASA) Only lithosphere has the strength and the brittle behavior to fracture in an earthquake.

The map below locates earthquakes around the globe. They are not evenly distributed; the boundaries between the plates grind against each other, producing most earthquakes. So the lines of earthquakes help define the plates:

(from the USGS)

In cross section, the Earth releases its internal heat by convecting, or boiling much like a pot of pudding on the stove. Hot asthenospheric mantle rises to the surface and spreads laterally, transporting oceans and continents as on a slow conveyor belt. The speed of this motion is a few centimeters per year, about as fast as your fingernails grow. The new lithosphere, created at the ocean spreading centers, cools as it ages and eventually becomes dense enough to sink back into the mantle. The subducted crust releases water to form volcanic island chains above, and after a few hundred million years will be heated and recycled back to the spreading centers.

Earthquake occurrence in different plate tectonic settings:

The map below of Earth’s solid surface shows many of the features caused by plate tectonics. The oceanic ridges are the asthenospheric spreading centers, creating new oceanic crust. Subduction zones appear as deep oceanic trenches. Most of the continental mountain belts occur where plates are pressing against one another. The white squares locate examples given here of the different tectonic and earthquake environments.
(topography from NOAA) There are three main plate tectonic environments: extensional, transform, and compressional. Plate boundaries in different localities are subject to different inter-plate stresses, producing these three types of earthquakes. Each type has its own special hazards.

At spreading ridges, or similar extensional boundaries, earthquakes are shallow, aligned strictly along the axis of spreading, and show an extensional mechanism. Earthquakes in extensional environments tend to be smaller than magnitude 8. (Click here for an explanation of earthquake magnitude).

A close-up topographic picture of the Juan de Fuca spreading ridge, offshore of the Pacific Northwest, shows the turned-up edges of the spreading center. As crust moves away from the ridge it cools and sinks. The lateral offsets in the ridge are joined by transform faults.

(from RIDGE, LDEO/Columbia Univ.)

A satellite view of the Sinai shows two arms of the Red Sea spreading ridge, exposed on land.

(from NASA)

Extensional ridges exist elsewhere in the solar system, although they never attain the globe-encircling extent the oceanic ridges have on Earth. This synthetic perspective of a large volcano on Venus is looking up the large rift on its flank.

(from NASA/JPL)

(from the USGS)

At transforms, earthquakes are shallow, running as deep as 25 km; mechanisms indicate strike-slip motion. Transforms tend to have earthquakes smaller than magnitude 8.5.

The San Andreas fault in California is a nearby example of a transform, separating the Pacific from the North American plate. At transforms the plates mostly slide past each other laterally, producing less sinking or lifing of the ground than extensional or compressional environments. The yellow dots below locate earthquakes along strands of this fault system in the San Francisco Bay area.

(from NASA/JSC; topography from NOAA)

At compressional boundaries, earthquakes are found in several settings ranging from the very near surface to several hundred kilometers depth, since the coldness of the subducting plate permits brittle failure down to as much as 700 km. Compressional boundaries host Earth’s largest quakes, with some events on subduction zones in Alaska and Chile having exceeded magnitude 9.

This oblique orbital view looking east over Indonesia shows the clouded tops of the chain of large volcanoes. The topography below shows the Indian plate, streaked by hotspot traces and healed transforms, subducting at the Javan Trench.

Sometimes continental sections of plates collide; both are too light for subduction to occur. The satellite image below shows the bent and rippled rock layers of the Zagros Mountains in southern Iran, where the Arabian plate is impacting the Iranian plate.

(from NASA/JSC)

Nevada has a complex plate-tectonic environment, dominated by a combination of extensional and transform motions. The Great Basin shares some features with the great Tibetan and Anatolian plateaus. All three have large areas of high elevation, and show varying amounts of rifting and extension distributed across the regions. This is unlike oceanic spreading centers, where rifting is concentrated narrowly along the plate boundary. The numerous north-south mountain ranges that dominate the landscape from Reno to Salt Lake City are the consequence of substantial east-west extension, in which the total extension may be as much as a factor of two over the past 20 million years.

(Topo map from the Lamont-Doherty Earth Observatory of Columbia Univ.; motions added from published GPS results.)
The extension seems to be most active at the eastern and western margins of the region, i.e. the mountain fronts running near Salt Lake City and Reno. The western Great Basin also has a significant component of shearing motion superimposed on this rifting. This is part of the Pacific - North America plate motion. The total motion is about 5 cm/year. Of this, about 4 cm/yr takes place on the San Andreas fault system near the California coast, and the remainder, about 1 cm/year, occurs east of the Sierra Nevada mountains, in a zone geologists know as the Walker Lane.

As a result, Nevada hosts hundreds of active extensional faults, and several significant transform fault zones as well. While not as actively or rapidly deforming as the plate boundary in California, Nevada has earthquakes over much larger areas. While some regions in California, such as the western Sierra Nevada, appear to be isolated from earthquake activity, earthquakes have occurred everywhere in Nevada.

J. Louie, 11 May 2001 (with contributions from J. Anderson)

More information About Earthquakes

source: http://www.seismo.unr.edu