Which earthquake waves have the lowest frequencies

The currently accepted method is the moment magnitude scale, which measures the total amount of energy released by the earthquake. At this time, seismologists have not found a reliable method for predicting earthquakes. A seismograph produces a graph-like representation of the seismic waves it receives and records them onto a seismogram.

Seismograms contain information that can be used to determine how strong an earthquake was, how long it lasted, and how far away it was. Modern seismometers record ground motions using electronic motion detectors. The data are then kept digitally on a computer.

If a seismogram records P-waves and surface waves but not S-waves, the seismograph was on the other side of the Earth from the earthquake because those waves cannot travel through the liquid core of the earth.

The amplitude of the waves can be used to determine the magnitude of the earthquake, which will be discussed in a later section. In order to locate an earthquake epicenter, scientists must first determine the epicenter distance from three different seismographs. The longer the time between the arrival of the P-wave and S-wave, the farther away is the epicenter. So the difference in the P and S wave arrival times determines the distance between the epicenter and a seismometer.

This animation shows how distance is determined using P, S, and surface waves. The scientist then draws a circle with a radius equal to the distance from the epicenter for that seismograph. The epicenter is somewhere along that circle. This is done for three locations. Using data from two seismographs, the two circles will intercept at two points. A third circle will intercept the other two circles at a single point. This point is the earthquake epicenter. Although useful for decades, this technique has been replaced by digital calculations.

The complexity of source rupture process or localized strong heterogeneities near a source fault might modify the radiation pattern of high-frequency P waves near the source region. Such characteristics suggest that the cause of these observations might be seismic wave scattering due to small-scale velocity heterogeneities along the propagation paths. To evaluate the effects of seismic wave scattering on the distortion of apparent P -wave radiation pattern, we conducted 3-D FDM simulations of seismic wave propagation using stochastic random velocity structure models.

These results might imply that scattering attenuation is not the dominant mechanism of P -wave attenuation in the crust of Chugoku region, southwestern Japan. Aki K. Attenuation of shear waves in the lithosphere for frequencies from 0. Earth planet. Google Scholar. Richards P. Quantitative Seismology 2nd edn University Science Books. Google Preview. Asano Y.

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Yabe S. Ide S. Seismic P-Wave Behavior. A reflected wave enters and exits at the same angle measured to the normal of the boundary - angle of incidence equals angle of reflection. If angle r equals 90 degrees, then the refracted wave propagates along the boundary interface. Simple Refraction Model. Seismic waves are generated from a source e. The reflected rays are always later than the direct ray. Its amplitude is stronger than the reflected ray, but is still later than the direct ray.

Two Layer Dipping Model. For a simple case of two dipping layers, seismic refraction can be utilized to calculated dip of the layers. Horizontal Multi-Layer Model. Interpretation of T-X plots remains the same. Problems and Limitations. Conformable sequences of sedimentary rock may form planar boundaries. However, erosion and uplift easily produce irregular boundary contacts.

More sophisticated algorithms can process refraction surveys where irregular interfaces might be expected. Typically, a profile can only detect features at a depth of one-fifth survey length. Thus, refraction imaging of the Moho would require profile lengths of over one hundred kilometers; an difficult experiment. In the hidden slow layer senario, a buried layer is overlain by a faster layer. No critical refraction will occur along the boundary interface.

Thus, refraction will not easily detect the slow layer. All is not lost since reflection seismology could detect the slower layer. If a thin layer produces first arrivals which cannot easily be identified on a seismogram, the layer may never be identified. Thus, another layer may be misinterpreted as incorporating the hidden layer. As a result, layer thicknesses may increase. Reflection seismology began to take prominence in the s to begin to locate salt domes, an indication where oil would be found.

The reflection method soon replaced the refraction after it was proved with numerous successes, the most visible in the petroleum industry. The key is to develop an equation which represents the time it takes for a particular ray to travel through this single layer. First, the seismic velocity through the layer of material that the wave is propagating needs to be lower than the layer directly below, which we will assume is infinitely thick.

Therefore, just by simple time-velocity relation and geometry:. This can be re-written dropping the subscripts as:. Well, the first thing to note is what you can do with the hyperbola. A hyperbola has an asymptote along which the hyperbola approaches. The equation of this line is.

Therefore, the asymptote for the travel time curve has a slope of the reciprocal of the velocity. Another approach to analysing the data is to get v elocity and thickness from a plot of x 2 v t 2. Now recall:. By squaring both sides, the equation resembles closely the equation of a straight line.

The slope of the line is the reciprocial of the square of the velocity. The intercepts gives h via:. Exploration Seismology. In the exploration industry there are many ways of processing reflection data so as to provide more information about the near sub-surface. This is beyond this course, but you may read more non-examinable material , and also in the following text taken from the Signalworks Pty.

Ltd web site. R eflection seismology is a technique for imaging the geological structure beneath the earth's surface using sound energy. The technique is used primarily for oil exploration. An acoustic energy source at the surface transmits an acoustic signal into the earth, which reflects some of the energy back toward the surface at each geological interface. An array of geophones or hydrophones detects the faint signals reflected back to the surface, which are recorded for later processing.

The raw data is very noisy and uninterpretable, requiring extensive processing to produce an image of the earth's interior. F igure 1 illustrates the process of marine seismic data acquisition.

The survey ship trails an acoustic source usually compressed air 'guns' and a string of hydrophones, called a streamer. The streamer is usually about m in length and contains groups of hydrophones spaced typically every 15m. When the air guns are fired, releasing a pulse of compressed air, a pressure pulse radiates in an approximately spherical wavefront through the water and into the earth. The semi-circles in figure 1 indicate the position the wavefront at regular intervals in time say every mS.

When the wavefront reaches a reflecting geological boundary, some of the wavefront energy is reflected back towards the surface light grey semi-circles. This echoed acoustic energy is sensed by the hydrophones and recorded on the ship for later processing. To simplify seismic acquisition models, the energy received at a hydrophone can be considered to have travelled along a linear raypath from the source, into the earth, then reflecting from the boundary back to the hydrophone.

Raypaths from the source to four hydrophones are shown in figure 1. The raypaths are perpendicular to the wavefronts. A coustic imaging in its simplest form consists of measuring the time taken by a pulse to travel from a source to a reflector and back to a receiver.

Repeating these measurements over a range of positions allows an image of the reflecting surface to be formed. Figure 2 shows the configuration of a simple imaging system. In practice, noise and imaging distortions require more elaborate data acquisition configurations and data processing techniques to achieve accurate imaging.