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:

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.

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.

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.

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.

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:

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):

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)

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.

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

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.

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.