Seismology

Short definition of seismologySeismology is one of the geophysical studies which examines the propagation of seismic wavesthrough the Earth.

Division of seismology:earthquake seismology,applied seismology.

Some important properties of earthquake seismology:seismic source with high energy,seismic source with random occurrence,very large investigation depth (> 10 km),the seismic waves are detected by seismometers,seismometers are located at seismological observatories,seismological observatories form a global network for listening the Earth's activity both in time

and space

Some important properties of applied seismology:man-made seismic sourcecontrolled application of the sourcethe energy of source depends on the required investigation depth,the seismic waves are detected by geophones (on the ground surface) or hydrophones (on the water surface),investigation depth ranges from one meter to a few kilometers,measurements are designed in advance.

Division of applied seismology:exploration seismology,near-surface, or shallow, seismology.

Some important properties of applied seismology:focuses on the study of deeper geologic structures, investigation depth typically varies between a few kilometers and a few hundred meters,frequently used in the exploration of raw materials.

Some important properties of shallow seismology:frequently used in the solution of hydrogeological, environmental and geotechnical problems,typical investigation depth ranges from a few hundred meters to one meter,

Most frequent applications of shallow seismology:location of water table, water bearing fractured zones, cavities, sinkholes,delineation of faults and fractures,determination of depth to bedrock.

Solid mechanics

Seismic investigations are based on the fact that seismic waves propagate with different

velocities in different rocks. In order to understand how seismic methods work and what we can expectfrom them, we must know the principles of solid mechanics.

The so-called elastic body is a theoretical model in solid mechanics. By the application of thismodel, we can describe the behavior of solid materials, especially their motion and deformation, whenthey are under the action of forces. The most important property of an elastic body is the elasticity.Elasticity is the ability of a body to resist a distorting influence and to recover its original size andshape when the influence is removed. Fluids do not show elastic behavior, because they have no fixed shape and are easily deformed. Butsolid rocks can be considered as elastic bodies in the range of very small deformations. Because thepropagation of seismic waves through rocks usually entails very small deformations, the elastic bodycan be considered as a good enough approximation of rocks.When an external force begin to act on the surface of an elastic body, an internal force will arise withinthe body. This internal force called stress is a surface force which acts across the internal surfaceelements of the body.

https://en.wikipedia.org/wiki/Stress_(mechanics)This figure demonstrates the stress across a surface element (yellow ellipse) separating two volumeelements of the body. The greater arrow located at the center of the surface element represents theresultant of the stress distributed over the surface element (small arrows).

Since the magnitude and the direction of stress can change from point to point within the body, thestress is expressed in the form of a vector field mathematically. This stress field describes the stressstate of a body.The dimension of stress is the same as that of pressure, so its unit is pascal (Pa, newton per squaremeter) in the International System, or psi (pound per square inch) in the Imperial system. But the magnitude of stress expressed in Pa is usually a very high value, so the use of MPa is morecommonly in engineering practice.The stress field causes the displacement of particles within the body which means the deformation ofthe body. The deformation is described by the so-called strain which gives the relative displacement ofparticles within the body. Relative displacement means that the final position of a particle is related to its initial position.Because the strain is the ratio of two quantities with the same unit, it is a dimensionless quantity.

https://www.nde-ed.org/EducationResources/CommunityCollege/Materials/Mechanical/StressStrain.htm

This figure shows a simple case of the deformation which occurs when a bar is stretched by tensionforces. The strain in this bar is the ratio of the change in length and the original length.

Similarly to the stress, the strain is also expressed in the form of a vector field mathematically. Thisstrain field describes the deformation state of a body.In the case of elastic deformation, a body completely recovers its original shape and size after the stressfield has been removed. It means that the elastic deformation is reversible. On the contrary, the plasticdeformation is irreversible.

Elastic stress and strain in engineering practice

Loading means the application of an external force or forces to an object. Depending on the type ofloading, different stress and strain fields arise within a body. There are five fundamental loading conditions:

tension, compression, bending, shear,and torsion.

From the seismic wave propagation point of view, tension, compression and shear are of importance.

A stress vector acting across a unit surface can be decomposed into two components:the normal stress (perpendicular to the unit surface), and the shearing stress (parallel to the unit surface).

http://homepage.ufp.pt/biblioteca/WebBasPrinTectonics/BasPrincTectonics/Page2.htmThe normal stress component (n or ) is perpendicular to the unit surface and it is also calledcompressive, or tensile stress, depending on the type of loading (compression or tension).

Both compression and tension forces act along a common axis on the opposite sides of a body, and thedirections of the forces are reversed. Compression causes the contraction of the body, while tensionresults in the extraction of the body along the axis of loading. Compressive stress (with a negative sign,n< 0) is generated by compression, and tensile stress (with a positive sign, n> 0) is generated bytension. Independently from its sign, normal stress always entails a deformation which changes boththe shape and the volume of the body.

http://magnet.fsu.edu/~odom/1000/deformation/def.htm

The figure above demonstrates the effects of tension and compression.

The other component of a stress vector is parallel to the unit surface and it is referred to as shear stress(dented by st or t). Shear stress is caused by the shearing component of an external force which acts onthe surface of a body tangentially. A shear stress field entails a deformation which changes only theshape of the body. The volume remains the same.

http://magnet.fsu.edu/~odom/1000/deformation/def.htmlThe figure above demonstrates the effects of shear.

Beside the stress fields introduced previously (compressional, tensional and shear), there is anotherstress field, the so-called volumetric stress (sV), which is also important in engineering practice. Avolumetric stress field occurs when the body is equally loaded by compressional forces in alldirections. Similarly to the normal stress, the volumetric stress entails a deformation which modifiesboth the shape and volume of a body.

http://magnet.fsu.edu/~odom/1000/deformation/def.htmlThe figure above demonstrates the loading causing the volumetric stress.

During the deformation caused by a loading, the displacement of a line segment between twoneighbouring points of a body can be divided into two components:

a change in length,an angle change.

The change in length is quantified by the so-called engineering strain, or Cauchy strain. It gives theratio of change in length to original length by the formula below:

where eng denotes the engineering strain, L is the original length of the line segment and DL representsthe change in length. Because this type of strain is calculated by the division of two quantities with thesame unit, it is a dimensionless quantity.In relation to compression or tension, we can often meet the terms longitudinal (or axial) andtransverse strains. Longitudinal strain is actually the engineering strain measured along the longeraxes of an elongated body, while transverse strain means the engineering strain measured in the cross-sectional direction of the body. They can be expressed by the following formulae:

where L is the original length of an elongated body along the axial direction,L is the change in length along the axial directionD is the original length of an elongated body along the cross-sectional direction,and D is the change in length along the cross-sectional direction.

http://physicscatalyst.com/mech/elasticity_2.phpThis figure demonstrates the meaning of quantitiesused for the definitions of longitudinal and transverse strains.

The angle change, which is the other effect of a deformation, is quantified by the angular distortionwhich gives the angle change of two originally perpendicular line segments within the body.

The angular distortion is measured in radian. The so-called shear strain is the tangent of the angulardistortion:

where is the angular distortion. Similarly to the previously introduced types of strain, the shear strainis also a dimensionless quantity.

https://en.wikipedia.org/wiki/Shear_modulusThis figure illustrates how the shear strain can be derived from the angular distortion g, the original

length of the line segment of a body L and the transverse displacement Dx.

The so-called volumetric strain is induced by a volumetric stress field and it is defined as the ratio ofchange in volume to the original volume:

where V denotes the original volume of a body and V is the volume change. It is a dimensionlessquantity.

http://magnet.fsu.edu/~odom/1000/deformation/def.htmlThe figure above demonstrates the volumetricdeformation entails a volumetric stress field.

Elastic moduli for homogeneous and isotropic materials

Stress and strain are strongly connected with each other because both of them are generated by thesame influence, which is mostly an external force. The relationship between the stress and strain isprovided by a property of matter called elastic modulus.Elastic modulus is a quantity which measures the resistance of a material to elastic deformation. So,the elasticity of a material is characterized by the value of elastic modulus.Depending on the type of stress and the measurement configuration, several types of elastic moduli aredefined. The most frequently used four of them are the following:

Young's modulus (or elastic modulus),Poisson's ratio,shear modulus,bulk modulus.

Young's modulus (often represented by E), or elastic modulus, measures the tensile elasticity of abody. In other words, it gives the tendency of a material body to deform along its length when tensileforces act on it. It is defined as the ratio of tensile stress to longitudinal strain.

where E is the Young's modulus,F is the tensile force,A is the cross-sectional area through which the force is applied,ΔL is the change in length,and L is the original length of the body.

The unit of the Young's modulus is pascal in System International.

http://hydrogen.physik.uni-wuppertal.de/hyperphysics/hyperphysics/hbase/permot3.htmlThe figureabove demonstrates the measurement configuration by which the Young's modulus of a material can be

determined.

In order to stretch a material body with a higher value of Young's modulus to a degree, a greater forcehas to be applied to it than to another one with lower value of Young's modulus. A soft material has alow value of Young's modulus, while a stiff or rigid material has a very high value of Young's modulus.Hence, we can say that the Young's modulus is the measure of stiffness.

Poisson's ratio (frequently denoted by n) is the measure of the so-called Poisson effect. Thisphenomenon means that a body compressed in one direction modifies their sizes in the other twodirections perpendicular to the direction of compression. Poisson's ratio is defined as the negative ratioof transverse strain to longitudinal strain.

where L is the original length in the direction of compression,DL is the change in length along the direction of compression,D is the original length in the transverse directionand DD is the change in transverse length.

Poisson's ratio is a dimensionless quantity.

http://www.pavementinteractive.org/article/poissons-ratio/The figure illustrates the measurementconfiguration by which the Poisson's ratio of a material can be determined.

Most materials expand in transverse directions when it is compressed in the axial direction. TheirPoisson's ratio is positive and ranges from 0 to 0.5. A perfectly incompressible material would have aPoisson's ratio of value 0.5. But in reality, there is no perfectly incompressible material.

Shear modulus, or modulus of rigidity, ( frequently denoted by G or m) measures the tendency of amaterial body to the deformation when a pair of shear forces acts on it. A pair of shear forces meansthat a force acts along one of the surface of a body and an opposing force with the same magnitude actsalong the opposite surface. The shear modulus is defined as the ratio of shear stress to shear strain.

where Fshear is the magnitude of the shear force,A is the area along which the force actsand g is the angular distortion cause by the force.

The unit of shear modulus is pascal in System International.

https://en.wikipedia.org/wiki/Shear_modulusThe figure illustrates the measurement configuration bywhich the shear modulus of a material can be determined.

A material with a higher value of shear modulus offers a stronger resistance to shearing deformation.Fluids have a shear modulus of 0 Pa, which means that they cannot resist shearing deformations at all.

Bulk modulus ( often represented by K) measures the volumetric elasticity of a material body, whichcharacterizes the tendency of a body to deform when it is equally compressed in all directions. It is

defined as the ratio of volumetric stress to volumetric strain.

where Fcompression is the compressional force acting on the surface of the body, A is the surface of the body, V is the original volume of the body, and V is the change in volume.

The unit of bulk modulus is pascal in System International.

http://faculty.uaeu.ac.ae/~maamar/physics1/121.htmlThe figure demonstrates the measurementconfiguration by which the bulk modulus of a material can be determined.

The lower the value of bulk modulus the more compressible the material. For example, gaseousmaterials are characterized by low values of bulk modulus. The inverse bulk modulus gives thecompressibility of a material ( often denoted by ). Actually, the bulk modulus can be considered asan extension of Young's modulus to three dimensions.

The values of elastic moduli can be considered as constant for each matter if the loading is static. Staticloading means that either the magnitude or the direction of loading do not vary with time. However, inthe case of dynamic loading, the value of an elastic modulus depends on not only the quality of matterbut the frequency of change in loading.

Wave theory

Elastic wave propagation is a spatial and temporal variation of the stress and strain fields within anelastic body. Rocks can be considered as elastic materials from the perspective of elastic wavepropagation.There is a very close interaction between stress and strain fields. A change in one of them causes thechange in the other, and vice versa.Since the seismic waves are actually low-frequency elastic waves, and the seismic methods are basedon the fact that seismic waves coming from a source are able to propagate through rocks, it is importantfor us to get acquainted with the basics of elastic wave theory.

Elastic waves

An elastic wave is a wave by which a disturbance in the stress-deformation state propagates through amedium both in space and time. The source of the disturbance can be an impulsive or a periodic changein the stress field, which occurs at a point or a relatively small part of a medium. Of course, a sudden change in the stress field is always induced by a change in the external forces

acting on the body. The perturbation in the stress field causes a change in the strain field, and thedeformation pattern corresponding to this change will propagate outward from the location of source asan elastic wave.It is important to note that not the particles but the change in the stress and strain fields travel throughthe medium during the propagation of an elastic wave. The particles are oscillating about theirequilibrium positions and the variation of the phase of oscillation both in space and time forms theelastic wave.

There are two principal types of elastic waves: body waves and surface waves.

A body wave is a wave which travels three-dimensionally through an elastic medium.A surface wave is a wave which propagates along and near by the surface of an elastic medium.

http://www.parkseismic.com/Whatisseismicwave.htmlThe figure tries to illustrate the differencesbetween body and surface waves from wave propagation point of view. Notice that the amplitude of

surface waves quickly decays with depth.

Body waves

On the basis of the deformation pattern propagating through the medium, two types of body wave canbe differentiated:

compressional wave, or P-wave (P comes from “primary”),and shear wave, or S-wave (S comes from “secondary”).

During compressional wave propagation, periodic alternations of contraction and expansion are takingplace within the medium. Compressional waves belong to the group of longitudinal waves (similarly tothe sound waves), because the particles are oscillating along axes parallel to the direction of wavepropagation. This type of wave motion entails periodic change both in volume and shape. Because thestress field has only a normal component, a compressional wave can propagate not merely in solids butalso in fluids.

http://www.geo.mtu.edu/UPSeis/waves.htmlThis figure illustrates the compressional wave propagation.The square prisms symbolize the same piece of an elastic medium at succeeding time moments (T=0, 1,

2, 3). The small cubes within the prisms represent the particles of the medium. It can be seen as thecompression and dilatation of particles travel in the direction of wave propagation (along the Y-axis).

During shear wave propagation, periodic alternations of shear deformation are taking place within themedium. Shear waves belong to the group of transverse waves (similarly to the light waves), becausethe particles are oscillating along axes perpendicular to the direction of wave propagation. Shear wavepropagation entails periodic change only in shape (there is no volume change). Because the stress fieldhas only a shear component, and fluids are not able to resist shear deformation, a shear wave canpropagate merely in solids.

http://www.geo.mtu.edu/UPSeis/waves.htmlThis figure illustrates the shear wave propagation in asimilar way than it was shown previously for the compressional wave. It can be seen that the motions

of the particles are perpendicular to the direction of wave propagation.

Velocity of body waves

The velocity of a body wave depends on the elastic properties and the density of the medium throughwhich it travels. As we could see previously, the elastic properties of a medium are characterized by theelastic moduli. Because different stresses and deformations are connected to the two types of bodywave, they propagate with different velocities in the same medium. The velocities of compressionaland shear waves can be expressed by the following formulae:

where VP is the velocity of compressional wave (P-wave), VS is the velocity of shear wave (S-wave), is the density of the medium,K is the bulk modulus of the mediumand is the shear modulus of the medium.

In the case of fluids (for example water or air), the value of shear modulus () is equal to zero. (Itmeans that fluids cannot resist shear deformations at all.) This is the reason why a compressional wavepropagates slower in fluids than in solids. Due to this effect of fluids on the wave velocity, the increaseof porosity results in the decrease of compressional wave velocity in rocks.There is another consequence of the zero value of shear modulus () in fluids. Namely, the shear wavevelocity is zero for all the fluids. That agrees with the fact that shear waves cannot propagate throughfluids.Since the value of bulk modulus (K) is positive, the compressional wave velocity (Vp) is always greaterthan the shear wave velocity (VS). So, a compressional wave is faster than a shear wave in the samesolid material. Due to this relationship, the compressional wave arrives first at a receiver located on the ground surfacefar enough from the source of the elastic waves. This is the reason why the other name of this wave isprimary, or P-wave. (If the distance between the source and receiver is not long enough, surface wavescan overtake the compressional wave, because the way has to be run by surface waves is shorter thanthose of body waves.)The second arrival will generally be the shear wave, whose other name is secondary, or S-wave. Fromthe point of view of seismic methods, mostly the first arrivals (or first breaks), that is the compressionalwaves, are of importance.

http://depthome.brooklyn.cuny.edu/geology/onlinecore/plates/platequiz.htmThis figure shows a full waveform seismic signal obtained by detecting the arrivals of different waveswith a seismic receiver located on the surface. We can see the usual order of seismic wave arrivals:

compressional (or P-wave), shear (or S-wave), and surface wave components. It can also be noticedthat a detected seismic signal always contains some random noise as an undesired component.

Surface waves

There are two main types of surface waves: Rayleigh wave (also called ground roll),and Love wave.

Surface waves are slower than body waves, so their arrivals usually follow the body waves on a fullwaveform seismic signal detected by a receiver located on the surface. Because the surface wavemotion strongly decays with depth, it is limited to the vicinity of ground surface.In comparison with body waves, the movement of particles is rather complex during the propagation ofsurface waves. The Rayleigh waves are characterized by elliptical retrograde particle motions in avertical plane.

http://depthome.brooklyn.cuny.edu/geology/onlinecore/plates/platequiz.htmThis figure try to illustratethe propagation of a Rayleigh wave both in time and space.

Love waves are horizontally polarized shear waves (SH waves), which means that the movement ofparticles are limited to horizontal planes and transverse to the direction of wave propagation. This typeof surface wave exists only in the case when the subsurface structure consists of at least two layers, andthe velocity of wave propagation is higher in the lower layer.

http://depthome.brooklyn.cuny.edu/geology/onlinecore/plates/platequiz.htmThis figure illustrates thepropagation of a Rayleigh wave both in time and space.

The effect of surface waves can be very destructive in case of earthquakes. From the point of view ofseismic methods, the surface waves are not useful. In fact, they are considered as noise and can causedifficulties in the recognition of compressional wave arrivals on a recorded seismic signal. Therefore, itis very important to eliminate or at least decrease the effect of surface waves on recorded seismicsignals by using signal processing techniques.

Direct and air waves

There are two other types of waves which are usually detected during a seismic measurement:direct waveand air wave.

Both of them are generated by the seismic source, but they do not provide any information about thesubsurface geology.

A Direct wave is a seismic body wave which travels through the upper layer directly from the source tothe receivers without meeting any subsurface layer boundary. Because of its short route, the directwave arrives sooner in the receivers than the body waves encountered subsurface layer boundaries.

An air wave is a wave which travels through the air directly from the source to the receivers. It is easyto recognize the arrival of air wave on a detected seismic signal, because it travels at a speed of 330m/s,(which is the speed of sound in air).

The figure try to demonstrate the routes of different waves coming from a seismic source.Some basic properties of waves

In order to review the most important properties of waves shortly, the selection of a simple sinefunction is good enough. The sine function is a periodic function which repeats over intervals of 2.This function is suitable to describe a physical quantity which varies periodically as a function of timeby the following form:

where t denotes the time (which is the independent variable of the function), u means the physicalquantity (which is the dependent variable of the function), A is the amplitude, T is the period, f is thefrequency, and is the phase.

http://www.doctronics.co.uk/signals.htmThe figure shows a sine function with its most importantproperties.

The amplitude (A) gives the maximum extent of the peaks in a wave. Its dimension and unit are thesame as that of the dependent variable.The period (T) gives the time interval between successive peaks (or troughs) in a wave. Its unit issecond (s).The frequency, more exactly the temporal frequency, (f) is the reciprocal of the period (T) and gives

the number of periods taking place during a unit time. Its unit is hertz (1 Hz=1/s) in SI.The phase, or phase angle, () gives the initial state of a periodic change. The whole cycle of thesubsequent states takes a length of time between periods. The value of the phase can range from 0 to2. A periodic change with zero phase means that the change starts at zero-crossing going positive. Theinitial time moment of an investigated process is generally chosen for zero.The so-called phase difference gives the difference between the initial states (or phase angles) of twoperiodic processes having the same frequency.Two waves are in-phase if their frequencies and initial states are the same. Otherwise, the waves areout-of-phase.

http://www.doctronics.co.uk/signals.htmThe figure represents two waves being in-phase.

http://www.doctronics.co.uk/signals.htmThe figure represents two waves being out-of-phase.

When a quantity periodically varies not only with time but also with distance, another parameter isrequired for the description of variation along the distance.This additional parameter is the wavelength often denoted by . It gives the distance betweensuccessive peaks (or troughs) in a wave. Its unit is meter (m).

The figure tries to illustrate a wave motion along time and distance.

The velocity of wave propagation (V) can be obtained as the product of frequency and wavelength:

The velocity of wave is a physical property of the medium through which the wave propagates. So, thevalue of wave velocity is constant for each material. Accordingly, increasing frequency entailsdecreasing wavelength and vice versa.

Frequency range of seismic waves

A seismic wave is a type of elastic wave which travels through the Earth's interior. In reality, seismicwaves cannot describe a single sine wave, because they are much more complex processes. We mustimagine a seismic wave as a superposition of no end of sine waves with different frequencies,amplitudes and phases. The superposition of infinite periodic components forms the resultant wave.Each periodic component participates in forming the wave in different measure. The measure of thisparticipation as a function of the frequency gives the so-called frequency spectrum of the wave.If we neglect those lower and higher frequency components whose contribution is very small, we get afrequency range (or band) which contains the essential components of a wave. The frequency range ofseismic waves encompasses a wide interval. The frequency range of seismic waves depends on the source and the medium through which theypropagate:

earthquakes generate seismic waves with a low frequency range of 0.01 Hz to 2 Hz, seismic waves generated by artificial sources can be characterized by a higher frequency range of 10 to about 100 Hz.

These values are valid for the detected seismic waves. The generated seismic waves have a broaderfrequency range, but the higher frequency components quickly decay during the propagation.

In fact, a seismic wave is not a periodic process. It can be considered as a wavelet which has finiteextent both in time and space. A wavelet is a wave-like oscillation with an amplitude that begins at zeroand decreases back to zero at the end.

http://www.lohninger.com/wavelet.htmlThe figure above shows a wavelet.

On the basis of the frequency spectrum, a dominant frequency of the seismic wave can be determined.The dominant frequency gives the frequency component which mostly participates in forming thewave. If the wave velocity of the actual material and the dominant frequency are known, the dominantwavelength can also be determined by the relationship introduced previously.

http://www.wavelet.org/tutorial/wbasic.htmA full waveform seismic signal, which contains theobservation of several seismic waves arrived at a receiver successively, is made up of several wavelets

belonging to the different seismic waves.

Attenuation of seismic waves

The energy of a seismic wave (E) is proportional to the square of its amplitude (E ~ A 2). It means that ahigh energy wave is characterized by a high amplitude and conversely. As a seismic wave is travellingfarther and farther from the source, its amplitude along with its energy is attenuating gradually.The loss of energy is the consequence of two effects:

geometrical spreadingand intrinsic attenuation.

The geometrical spreading is actually a geometrical effect without any physical mechanism. Theessence of this effect can be summarized as follows.When a seismic wave generated by a point source propagates through a homogeneous medium, thepoints being the same phase of oscillation at a time moment form spherical wave fronts.The total energy of the wave spreads over these wave fronts. Each wave front has its own constantenergy. As the wave is getting farther and farther from the source, the surface of each wave front isbecoming larger and larger. Due to the increase of a wave front, the surface energy density (whichmeans the energy per unit area) is gradually decreasing. This is the geometrical reason why the energyof a seismic wave decreases with distance in all directions.

http://www.performing-musician.com/pm/apr09/articles/technotes.htm?print=yesThis figure shows

how the surface of a wave front increases with the distance from the source.

The intrinsic attenuation, which is the other type of loss of energy, is purely a physical effect. Owingto the displacements of particles in a wave motion, frictional dissipation is taking place continuously,which converts some part of elastic energy into heat.

These two effects (geometrical spreading and intrinsic attenuation) jointly responsible for theattenuation of wave amplitude. The measure of attenuation as a function of distance from the sourcecan be described for a homogeneous medium by the following relationship:

where A is the amplitude at distance r from the source, A0 is the initial amplitude, and is theabsorption coefficient of the material.The higher the value of absorption coefficient the larger the attenuation of amplitude and the absorptionof seismic energy. The value of absorption coefficient in rocks ranges from 0.2 to 0.75 dB/wavelength.It depends on the frequency of wave and the lithological character of the rock. The more consolidated arock the smaller the value of its absorption coefficient.An ideally elastic medium has an absorption coefficient of zero. It means that there is no intrinsicattenuation within the medium. But in reality, there is no ideally elastic rock, so we always have to takeinto account the attenuation of amplitude.In the case of near-surface sedimentary structures, the degree of consolidation is generally low. Itmeans that larger intrinsic attenuation of amplitude can be expected for shallow seismic surveys thanfor deeper seismic surveys. (If the attenuation is too strong, the amplitude of a seismic wave cannot beseparated from the background noise on a detected signal. In such cases, the useful information contentof a seismic wave arrival is lost.)The higher the frequency of a wave the higher the value of absorption coefficient. It supports theempirically proven fact that higher-frequency waves attenuate more rapidly than lower-frequency ones.Due to this effect, the higher frequency components gradually disappear in a seismic wave as it ispropagating through rocks. Thus, the seismic waves penetrated the deeper parts of a subsurfacestructure contain rather lower-frequency components than higher-frequency ones. (Later we will seethat the disappearance of high frequency components has an unfavourable consequence for the verticalresolution of seismic reflection method.)

Ray theory

If we want to study in what directions the waves travel and how the boundaries modify the direction ofwave propagation, it is worth applying the approach of ray theory. Contrary to the wave theory, whichis suitable for the quantitative description of wave propagation, the ray theory provides a qualitativeoutlook by which we can trace the routes of waves in a layered subsurface structure (or half-space).Ray theory is generally used in geometrical optics where a ray is the idealized model of a light wave.The principles used for light waves are also applicable to seismic waves. So, the ray (or ray path) of aseismic wave is considered as a line which is perpendicular to the wave fronts along its trace andindicate the direction of wave propagation. This line also represents the path of energy flow connectingto the wave propagation.

https://1f308d6acfa3dcbec8dfb7385adde1ece594768d.googledrive.com/host/0B6TvZfgdBGQ8Mzg2MC1nSFpYazQ/dissertation/2%20Methodology.htmlThis figure tries to demonstrate the meaning of ray

path in seismic wave propagation. It can be seen that each ray path coming from a point source is a linewhich is perpendicular to the wave fronts along its trace.

http://www.crewes.org/ResearchLinks/Converted_Waves/Page2.htmlThe figure above shows a so-called horizontally layered half-space which is a frequently used geophysical model of sedimentary

basins. Some possible ray paths of seismic waves are also presented. Each of these ray paths starts froma seismic source on the surface, penetrates some of the layers, reflects from one of the layer boundaries,then arrives in a seismic receiver located on the surface. The travel-time values of the waves belonging

to different ray paths bear the information about the subsurface structure.

Reflection and refraction of seismic waves

When a seismic body wave (either a compressional or a shear wave) arrives at a layer boundaryseparating two media with different elastic properties, the energy of the wave is partly reflected fromthe boundary and partly transmitted to the second medium.Reflection means the phenomenon when a wavefront arrives at an interface between two differentmedia, the direction of propagation changes, and the wave front with attenuated energy returns into themedium from which it arrived.Refraction means the phenomenon when a wavefront arrives at an interface between two differentmedia, the direction of propagation changes, the wave front with attenuated energy passes the interfaceand penetrates the other medium.

In general, when an incident seismic wave encounters a layer boundary, it is divided into two reflected

waves and two refracted waves. One of the reflected waves is compressional and the other is shear.That is also true for the refracted waves.So, the following waves are generated at a layer boundary:

a reflected compressional wave (P-wave),a reflected shear wave (S-wave),a refracted compressional wave (P-wave),and a refracted shear wave (S-wave).

The division of an incident wave at a layer boundary involves that the energy of incident wave is alsodivided among the waves generated by the reflection and refraction. After occurring the reflection andrefraction, each of these waves travels its own trace in different directions.

Prem V. Sharma: Environmental and engineering geophysics, Cambridge University PressThis figureillustrates what happens when an incident compressional wave arrives at a layer boundary with anangle of incidence ip. The meanings of denotations appearing in the figure are the following: S in a

subscript is a shear wave, P in a subscript is a compressional wave, R is the angle of reflection, r is theangle of refraction, V is the seismic wave velocity, is the density of a layer, 1 in a subscript is the

layer 1, and 2 in a subscript is layer 2. The angles are defined relative to the axis of incidence, which isperpendicular to the layer boundary.

Snell's law

There is a relationship between the angles and velocities of the waves generating at a layer boundary. Itis given by the so-called generalized Snell's law.

By this law, the ratio of the sine of angle to the appropriate wave velocity is the same for each type ofwaves. This ratio gives a constant value, which is determined by the properties of incident wave (theangle of incidence and the velocity of incident wave in the upper layer). The angles of reflections andrefractions can be calculated by using this constant value as well as the wave velocities.

A special case of the reflection is the normal, or vertical, incidence. It occurs when the ray path of anincident compressional wave is perpendicular to the layer boundary (that is the angle of incidence iszero). In such a case, neither reflected nor refracted shear waves are generated. The total energy of theincident wave is split between the reflected and refracted (or transmitted) compressional waves.

http://www.ukm.my/rahim/Seismic%20Refraction%20Surveying.htmThe figure illustrates the case ofnormal incident.

In practice, we are not able to produce a measurement belonging to a normal incident, because aseismic receiver cannot be put on a seismic source without any damage. Only the near-normal, ornear-vertical, incidence can be implemented when the horizontal distance between the seismic sourceand the receiver is much smaller than the depth of layer boundary (that is the angle of incidence is asmall value). In this case, also the reflected and refracted compressional waves have determinant roles,because they convey the major part of original energy. This is the reason why the compressionalreflected and refracted waves are mostly used for mapping the subsurface structures in seismicprospecting.

http://www.ipims.com/data/gp13/P0535.asp?UserID=&Code=3776The figure illustrates the case ofnear-normal incident.

Critical refractions

There is a special case of refracted waves, as well. When the seismic wave travels faster in the lowerlayer than in the upper one (V2 > V1), which is often fulfilled in practice, the angle of refraction will begrater than the angle of incidence (r > i). This relation is the consequence of Snell's law.

It is also results from the Snell's law that increase in the angle of incidence entails increase in the angleof refraction. When the angle of incidence reaches a certain value, the angle of refraction will be

perpendicular to the axis of incidence. In such a case, the refracted wave will travel collaterally withthe layer boundary. This effect is called critical refraction, and the angle of incidence by which itoccurs is referred to as critical angle.The critical angle can be expressed from the Snell's law as follows:

The critically refracted compressional wave propagates with the velocity of the lower layer (VP2).During its propagation, the points of the lower layer located near by the boundary behave as sources ofsecondary compressional waves. These secondary waves will travel upwards through the upper layer,and their ray paths subtend critical angle to the normal of the interface. The superposition of thesecondary compressional waves (or P-waves) will form a plain wave front called head wave.

http://www.ukm.my/rahim/Seismic%20Refraction%20Surveying.htmThe figure demonstrates the caseof critical refraction.

By detecting the arrivals of the head wave front at different points of the ground surface, we can obtainarrival time data which contain information about the depth of layer boundary. The deeper the layerboundary the longer the arrival time values of the head wave. Of course, the measured arrival time dataare also influenced by the wave velocities in the layers. Actually, the seismic refraction method isbased on the theory of critically refracted waves.

Reflection coefficient

Snell's law is a geometrical relationship which does not give any information about the relations of theamplitudes belonging to the different types of waves. In order to investigate the relations of amplitudes,we have to introduce the so-called acoustic impedance. The acoustic impedance is an acoustic property of the medium, which can be calculated by theproduct of density () and wave velocity (V):

I = ·VThe acoustic impedance of a given type of waves can be obtained by substituting the proper wavevelocity in the formula above. (For example, the acoustic impedance of an incident compressionalwave is calculated by using the velocity of compressional wave in the upper layer.)In the case of normal incidence (when the angle of incidence is zero), the ratio of the amplitude ofreflected compressional wave to the amplitude of incident compressional waves is characterized by theso-called reflection coefficient (R ).The reflection coefficient (R ) can be expressed by the following equation:

where AR is the amplitude of reflected compressional wave, Ai is the amplitude of incidentcompressional wave, I1 is the acoustic impedance of the compressional wave for the upper layer, and I2

is the acoustic impedance of the compressional wave for the lower layer.Since the density and the wave velocity generally increase with depth (due to the compaction andconsolidation) the value of reflection coefficient is mostly positive. A positive reflection coefficientmeans that there is no phase reversal between the reflected and incident waves. (Phase reversal meansthat the phase of the wave is shifted by radian (or 180°).The value of reflection coefficient for the interfaces between rock layers is usually small. It is rarelygreater than 0.2 (or 20%). Since the ratio of reflected energy to the incident energy is proportional tothe square of the reflection coefficient, its vale is frequently smaller than 0.01 (or 1%). It means thatonly a smaller part of the incident wave energy is reflected from a layer boundary. Yet, there are somespecial cases when this ratio can reach 70 % or even 100%. Such excellent reflectors are the surface ofthe oceans and the surface of the Earth itself (they are characterized by significant density contrast).

Refraction coefficients

Similarly to the reflection coefficient, another quantity is defined for the case of refraction. Therefraction coefficient (T) of an interface between two media gives the ratio of the amplitude ofrefracted compressional wave to the amplitude of incident compressional wave. It can be expressed bymeans of the acoustic impedance in the following way:

where AT is the amplitude of refracted compressional wave, Ai is the amplitude of incidentcompressional wave, R is the reflection coefficient, I1 is the acoustic impedance of the compressionalwave for the upper layer, and I2 is the acoustic impedance of compressional wave for the lower layer.The formula above is valid for the case of normal incidence.Both the reflected and refracted coefficients depend on the contrast between the acoustic impedances ofupper and lower layers. Since the density of rocks varies over a relatively narrow range, mainly thewave velocity of rocks determines the value of acoustic impedance. This is the reason why the velocitycontrast of two neighbouring media is the main factor influencing the energy relation between thereflected and refracted waves.

Diffraction of seismic waves

When a discontinuity, that is a sudden change, can be found along an interface separating two solidmedia (e. g. a faulted structure), the incident wave is diffracted. Due to the diffraction, thediscontinuity of an interface will serve as a source of scattered waves. The scattered waves will travelfrom this source and traverse the upper medium. During their propagation, they will meet other wavesreflected by the continuous segments of the interface.

Prem V. Sharma: Environmental and engineering geophysics, Cambridge University PressThe figureillustrates the diffraction of seismic waves

The meeting of diffracted (or scattered) and reflected waves entails their superposition. A net waveformgenerated by the superposition is the sum of individual waves. The scattered waves have a distortioneffect on the other reflected waves. Because of the distortion, the seismic image (called seismic section)of the subsurface structure will be blurred in the surrounding area of the discontinuity. This is thereason why mapping of layer boundaries is problematic near discontinuities.

http://www.xsgeo.com/course/mig.htmThe figure shows the effect of diffraction near a discontinuity of a reflecting horizon.

Seismic velocities in rocks

In order that a compressional wave can travel the same distance through different types of rocks,different time intervals are needed. This is because the composition, the structure and the texture of arock significantly influence the velocity of a seismic wave. From the point of view of seismic methods,

the most important property of a rock is the velocity at which the compressional wave propagatesthrough its material.The velocity of compressional wave depends on the density and the elastic properties (elasticmoduli) of a rock as we have seen it previously:

On the basis of the formula above, we could think that density decreases the wave velocity. But wemust not forget that elastic moduli also depend on density. Increasing density results in faster thanlinear increase of elastic moduli. So, the density has not only a direct but also an indirect effect on thewave velocity. As a result, the velocity of compressional wave increases with the density of rock.Since the increase of porosity reduces the bulk density of a rock, it decreases the velocity ofcompressional wave. The types of fluids filling the pore volume and the saturation for each fluidphase also influence the velocity of compressional wave. A fluid phase with increasing density andsaturation raises the magnitude of velocity. Hence, the velocity of compressional wave is greater in arock with water filled porosity than in a rock with dry (air-filled) porosity.The table below shows the velocity values of compressional wave for some materials and rocks. Someof them, which frequently occur in sedimentary basins, are highlighted by red colour.

Rock or material Vp (m/s)

air 330

water 1400-1500

ice 3000-4000

permafrost 3500-4000

weathered layer 250-1000

alluvium, sand (dry) 300-1000

sand (water saturated) 1200-1900

clays 1100-2500

glacial moraine 1500-2600

coal 1400-1600

sandstones 2000-4500

slates and shales 2400-5000

marls 2000-3000

limestones and dolomites 3400-6000

anhydrite 4500-5800

rock salt (halite) 4000-5500

granite and gneiss 5000-6200

basalt flow top (highly fractured)

2500-3800

basalt 5500-6300

gabbro 6400-6800

dunite 7500-8400

Prem V. Sharma: Environmental and engineering geophysics, Cambridge University Press

The velocity of compressional wave in rocks varies in a wider range than the density of rocks. Thehighest values of wave velocity are connected to the tight igneous and high grade metamorphic rocks(e.g. granite, gneiss ). But fracturing of these rocks decreases the velocity (see basalt flow top). In the case of sedimentary rocks, there can be significant differences between the values of wavevelocity even within the same formation. Compact and consolidated sedimentary rocks with lessporosity and higher bulk density provide faster propagation for the compressional waves. The lowestwave velocities, which are less than the wave velocity in water (< 1400-1500 m/s), occur inunconsolidated sediments with dry porosity nearby the surface (see dry sand and weathered layer inthe table). The zone of these sediments ranging from the surface to the water table is called the“weathered layer” or the low-velocity layer (LVL).

Field techniques in seismic surveys

In order to collect field data during a seismic survey, two main requirements have to be fulfilled.Firstly, seismic waves are to be generated on the ground surface by using a seismic source. Then, thetime intervals (travel-time) required for the seismic waves to travel from the source to different pointsof the surface should be measured.Depending on whether reflected or refracted waves are used for getting information about thesubsurface structure, we can differentiate two major methods:

the seismic reflection method and the seismic refraction method.

Independently of the method, the following components are required for implementing a seismic fieldmeasurement:

a seismic source,several seismic receivers (called geophones in seismic surveys),a seismograph,cables.

The main functions of these components can be summarized as follows. A seismic source is used togenerate seismic waves which propagate through rocks. The detection of ground motions caused byseismic waves and the conversion of these motions to an electric voltage signal are the tasks of ageophone. Collecting the signals coming from a series of geophones simultaneously, recording thesignals as digitized data sets, and other important operations are provided by the seismograph. Thesignal transmission between the geophones and the seismograph is supported by cables.

Signal and noise

From the perspective of geophysical measurements, it is very important to realize the meaning of signaland noise. In general, a signal is the variation of a quantity recorded as a function of time and/or space. A signal always contains some information about the phenomenon which induces the variation of thequantity.In science and engineering, a recorded signal is always noisy. A noisy signal is considered as the sumof a desired or useful component and an undesired component called noise. The desired componentbears the information in which we are interested.

From a noise correction point of view, two types of noise can be distinguished. The so-calledsystematic noise is the result of a deterministic process, and its effect can be corrected if the processcausing the noise is well-known and described by mathematically. On the contrary, the random noiseis caused by a stochastic process or processes, and can only be reduced (not eliminated completely) byusing signal processing techniques.In seismic surveys, the output of a receiver (or geophone) is a noisy voltage signal, which isproportional to the velocity or acceleration variations (depending on the type of geophone) of groundmotion at a receiver site. The so-called signal-to-noise ratio (abbreviated SNR or S/N) is a quantityused for comparing the level of a useful signal component to the level of background noise component.It is defined as the ratio of the useful signal power to the noise power. Its value is generally expressedin decibel (dB), which is based on a logarithmic scale as follows:

where Psignal and Asignal are the power and the amplitude of useful signal component, Pnoise and Anoise arethe power and the amplitude of noise (the power of signal is proportional to the square of signalamplitude). If the signal-to-noise ratio is grater than 0 dB, the rate of useful component in a signal ishigher than the rate of noise.

Seismic source

A man-made seismic source is a device or a material which is used to generate controlled seismicenergy on the ground surface. The technique of its application is closely connected with any type ofseismic source. The proper application is the fundamental condition of successful and safe seismicwave generation.For near-surface seismic surveys, the investigation depth is usually less than a kilometer. This is thereason why there is no need to transmit too high energy to the subsurface media. Consideration of the following factors are suggested for choosing the suitable seismic source:

cost,repeatability,convenience,efficiency,safety.

Repeatability (why is it important in some cases)

When the singular use of a source device is not able to provide a high enough seismic energy, therecorded signal will be too noisy at the sites of receivers. It means that the effect of seismic waves inthe recorded signal is strongly distorted by the background noise.In this case, the repetition of seismic wave generation and detection can be useful for the improvementof signal-to-noise ratio. During the repetitions, the sites of source and receivers are fixed.After the repetitions, the recorded signals of each receiver are summarized. In such a way, we get aresultant signal (stacked signal) for each receiver, which is less noisy than any recorded signal, becausethe summation is able to attenuate the effect of random noise and enhance the effect of seismic waves.This signal processing technique is called vertical stacking, and its result is known as vertical stackedsignal.

Penetration depth

A seismic source can be considered as a point source, because the size of area or volume contacted tothe source directly is much smaller then the wavelengths of generated seismic waves. In practice, thereare several sources used in a wide variety of environments.So as to select an adequate seismic source, we always have to find an acceptable compromise betweenthe penetration depth and the vertical resolution. The penetration depth or investigation depth givesthe vertical depth of the deepest layer boundary which can be detected in a given situation. It dependson several factors, among other things

the power of source, the frequency band of generated seismic waves, the geologic structure as well as the acoustic properties of rocks etc.

As it was mentioned formerly, the high-frequency seismic waves quickly attenuate with depth, so thedeeper structures can be revealed by seismic waves with lower dominant frequency. But the decrease offrequency has an undesired effect on the vertical resolution of seismic measurements.

Vertical resolution

The vertical resolution of a seismic measurement characterizes the ability to distinguish subsurfaceobjects vertically by means of recorded seismic signals. Similarly to the penetration depth, the verticalresolution also depends on several factors.The theoretical upper limit of the vertical resolution is basically determined by the dominantwavelength of a seismic wave. By this limitation, the vertical resolution cannot be better than thequarter of the wavelength:

where hmin is the minimum layer thickness which can be distinguished yet by a seismic wave withwavelength .As we could see earlier, the wavelength depends on the frequency and the wave velocity. Since thewave velocity in rocks cannot be modify, the refinement of vertical resolution (by decreasing thewavelength) necessarily entails the increase of frequency. But seismic waves with higher frequency arenot able to get into the deeper part of subsurface. This is the reason why we cannot reach highresolution along with deep penetration.A realistic estimation for the maximum achievable frequency in the near-surface seismic application isabout 500 Hz. If we consider several hundreds of meters per second (n x 100 m/s) as the velocity ofcompressional wave, the wavelength is limited to 2 m. It means half a meter (2 m/4 = 0.5 m) verticalresolution. But that would be extremely good. In practice, the resolution like this cannot be reached.

Types of seismic sources

There are two main types of sources in shallow seismic surveys:impulse type sources,vibration type sources.

An impulse type source is characterized by a very short time interval of wave generation. So, its effectcan be considered as instantaneous. The group of impulse type sources can be divided into further threesubgroups:

impact sources,

impulsive or shooting sources,explosion sources.

In the case of impact sources, the seismic waves are generated by the impact of a heavy, solid body onthe surface. One of the most frequently used impact sources is striking a steel plate on the ground by asledge hammer and the other is dropping a spherical weight onto the ground. The advantages of using a sledge hammer or a weight dropper device are the low cost and therelatively good repeatability. But they are not able to provide higher seismic energy, so the depth ofpenetration is less than a hundred meter. The impact sources usually produce seismic waves with adominant frequency ranging from 50-200 Hz

http://www.aksgeoscience.com/seismic.htmlThis photo demonstrate a sledge hammer with a steel plate.A cable is led into the head of hammer. When the plate is struck by the hammer, an electric circuit is

closed for a moment. Through this circuit, a trigger impulse is transmitted to the seismograph in orderto start data recording. In such a way, data recording can be synchronized with the time moment of

wave generation.

http://www.geoexpert.ch/equipment.htmlAn accelerated weight dropping device can be seen in thephoto. The acceleration implemented by a spring increases the energy of source. The weight can be

lifted by an electric motor or manually. Because of the small total weight, it is easy to move under field

circumstances.

http://www.geoexpert.ch/equipment.htmlThis is another type of weight dropper which is mounted on atruck. It provides a higher seismic energy due to the free fall weight dropping from a higher level with

a heavier weight.

The working principle of impulsive or shooting sources is the same as that of a rifle or a shotgun. Thedetonation of a small explosive charge in a metal pipe shoots a bullet or a wadding into the ground. Because shooting is a dangerous action, we must strictly observe the safety rules concerning the use ofthese devices. In most countries, such as Hungary, special permission from the authorities is needed forusing them in seismic surveys. Shooting devices are characterized by high repeatability and a penetration depth ranging from less thana hundred meter to a few hundred meters. The penetration depth depends on the soil conditions, theamount of charge, and the caliber of the shooting device. Firing into a water-filled shallow hole drilled previously is often used, because it provides nearly thesame effects during repetitions. The dominant frequency of generated seismic waves ranges from 80 to200 Hz.

http://firearmshistory.blogspot.hu/2011/03/utility-firearms-geology-and-shotguns.htmlThis photoshows a shotgun used for seismic wave generation. Such a shotgun is also known as "Betsy Gun".

Sources like this are planned for only seismic wave generation, so they cannot be used as a regular rifleor a shotgun. (The closed cylindrical end of the device provides protection against the bullets

rebounding from the ground surface.)

During an explosion, a significant amount of energy is released suddenly in a small volume whichgenerates a pressure wave. This pressure wave takes an impulse-like effect on the ground and generateseismic waves nearby the source. The amount of seismic energy provided by an explosive source depends on the type and the amount ofexplosive material. It can be high enough to support a penetration depth of a kilometer. But theapplication of explosive materials is the most dangerous of all the seismic sources. In addition, it is rather expensive, as well as the storage, the treatment and the application of explosivematerials are strictly regularized. Therefore special permissions from the authorities are needed to getsimilarly to shooting devices.The explosive devices can be divided into two classes by their sensitivity and power:

primary explosivesand secondary explosives.

The primary explosives (e. g. blasting caps) are very sensitive which means that a relatively smallinitial energy is enough to induce the detonation. Since primary explosives provide a relatively lowenergy, they can be used for investigating shallower subsurface objects (located less than a hundredmeter) or initializing secondary explosives.In seismic practice, the electric type blasting caps are the most frequently used primary explosives. Anelectric blasting cap is fired by a current impulse coming from a blasting machine located far from thesite of detonation. The blasting machine is connected with the blasting cap by a pair of long wires toensure the safe application. A single blasting cap provides the dominant seismic energy in the range of100-200 Hz.

https://en.wikipedia.org/wiki/Blasting_capThe photo demonstrates electric type blasting caps.

Another important use of primary explosives is to detonate more powerful and less sensitive secondaryexplosives such as TNT, dynamite, or plastic explosive. The explosive charges are usually fired in ashallow hole filled with water. It is worth drilling shallow holes for the shots, because the explosion ismore effective and less dangerous in a hole. Since a small amounts of explosive is generally enough for a shot (less than 10 gramms), a single holecan be used more times. Secondary explosives produce the dominant seismic energy in the range of 50-200 Hz. A greatdrawback of explosive sources that it is not possible to use them near roads and urban areas. But inrural areas with high water table an explosive source may be the better choice

https://www.flickr.com/photos/57768042@N00/2805186128/An explosion of a seismic charge can beseen here. The muddy water is ejected high from the hole by the energy of explosion.

Vibration type sources are able to generate seismic energy over an extended period (up to twentyseconds). The frequency of vibration continuously varies as a function of time. The wave train of avibration is called “sweep”. In the case of the most frequently used linear sweep, the frequency changes

linearly from the lowest frequency to the uppermost one (upsweep). When the seismic wave generation is implemented by using a vibration type source, the first step ofdata processing is eliminating the effect of vibration source from the raw data. This operation is basedon the cross-correlation of the source and recorded signals. This type of seismic wave generation alongwith the special data processing workflow is called Vibroseis technique.The main advantage of vibration type sources is that they are less dangerous for roads and urban ares.Another advantage of using them is that the recorded data is less sensitive to background noise (e.g.footsteps, traffic, single car). Vibration sources are rather expensive, because the vibrator unit is often mounted on a vehicle. Bylowering the base of the vibrator unit, physical contact is made between the ground and the unit. Thevibration process is controlled from the driver's seat via a computer.

http://www.apigeophysical.com/T2500.pdfA mini vibrator used for shallow seismic surveys can beseen here. The vibrator unit is mounted on a small truck and moved by a hydraulic arm. It provides a

vibration frequency range of 10-500 Hz.

http://www.velseis.com/velseis/history/The so-called Mini-Sosie means an alternative vibration typeseismic source and a data processing workflow belonging to it. Actually, a vibratory rammer is used for

generating seismic waves.It is effective in imaging near-surface structures. This source is less expensive, portable and applicable

in both rural and urban environments.

Shot

In seismic practice, the term “shot” means the event of seismic wave generation independently fromthe type of source. A given shot is identified by the time and location of its occurrence. Shot points arearranged along seismic lines.