Seismology (Greek seismos = earthquake + Greek logos = science) is a geophysical discipline that studies the formation of earthquakes, the propagation of seismic waves, their effect on the Earth’s surface and the objects on it, and the structure of the Earth itself. Some of the basic problems and questions dealt with by seismology relate to the processes that lead to the occurrence of earthquakes, the processes that take place during earthquakes, and the phenomena that are their consequence. Seismology is used in construction (earthquake engineering) and urban planning, in the assessment of hazards and risks from earthquakes, and in the exploration of gas and oil deposits, that is, in general, in determining the structure of the Earth’s interior. Professional work in seismology includes recording earthquakes, locating and cataloguing them, mapping effects, exchanging data with international institutions, etc.
Seismics is a set of geophysical methods based on the study of records of artificially induced earthquakes for the purpose of researching the geological structure of the Earth or another planet.
An earthquake is a sudden release of accumulated energy within a limited area on Earth. Today’s seismology knows several different causes and types of earthquakes, but the most common are tectonic earthquakes (> 90%) that occur on faults and are the result of tectonic forces. Released energy can be gravitational (collapse earthquakes), hydrodynamic (volcanic earthquakes), elastic (tectonic earthquakes), etc., so we distinguish volcanic (7%), collapse (3%), impact (e.g. meteorite impact), and artificial earthquakes (e.g. induced by nuclear explosions, filling and emptying of reservoirs, etc.). Natural earthquakes are quasi-periodic, meaning they recur at irregular intervals. Travelling elastic waves, i.e., seismic or earthquake waves, spread in all directions from the focus of the earthquake.
The place where an earthquake originates in the Earth’s interior is called the focus or hypocentre, and the point on the Earth’s surface vertically above is called the epicentre. According to the depth at which the focus of the earthquake is located, there are shallow (up to 60 km), medium deep (from 60 to 350 km) and deep earthquakes (depths greater than 350 km).
Faults are shear cracks in the rocks of the Earth’s crust through which the visible movement of rock blocks, or fault wings, has been achieved – the fault plane separates two adjacent fault wings, among which, in the case of inclined faults, the following are distinguished: the underwing wing (part of the rocks below the fault plane) and the roof wing (part of the rocks above the fault surface).
Left: Sketch of the fault: the grey surface indicates the fault surface between two fault wings: above the surface is the roof fault, and below the surface is the lifting fault wing, and the arrow indicates the displacement along the fault. Right: Sketch of the relationship between the fault, the focus of the earthquake, the epicentre and the seismological station
A tectonic earthquake is an earthquake that occurs due to the release of elastic deformation energy. The connection between the energy of elastic deformation and the occurrence of earthquakes is given by the so-called elastic rebound theory. The theory was put forward in 1911 by H.F. Reid based on the study of geodetic measurements along the San Andreas fault before and after the great earthquake in San Francisco that occurred on April 18, 1906. In his theory, Reid states that the fracture of rock, which causes a tectonic earthquake, is the result of elastic tension (stress) greater than the strength of the rock can bear. Exceeding the strength causes relative displacement along fault wings. The maxima of relative displacements are not reached suddenly with the onset of failure, but gradually over a longer or shorter period of time. Seismic waves are generated on the fracture surface.
Schematic representation of the gradual exceeding of the strength of the rock, which leads to fracture and earthquakes
Since there is a possibility that the total movement along the fault wings does not happen all at once, but in several steps, any sudden stop and start of movement will cause the formation of waves. The energy released during the earthquake was stored in the form of elastic deformation energy just before the beginning of the fracture.
The earthquakes that follow the main and often strongest earthquake are called aftershocks: they are usually weaker than the main earthquake and are unevenly distributed in time. As a rule, the stronger the main earthquake, the stronger and more frequent the aftershocks, and the longer the post-seismic period of increased seismic activity. The end of an earthquake series can only be declared afterwards when it is established that the seismic activity has weakened to a level corresponding to that before the main earthquake. Sometimes the main earthquake is preceded by weaker earthquakes – these are called foreshocks, but whether something is a foreshock can only be determined later, after an even stronger earthquake has followed in a certain time and a certain area.
Seismic waves are actually elastic waves that arise on the surface of a fracture on a fault plane or due to some disturbance (eg. a meteorite impact or an explosion). During an earthquake, about 90% of the total energy will be spent on friction and various thermal effects, while 10% will be released in the form of seismic waves. There are two basic groups of seismic waves: body waves and surface waves.
Body waves are created in the focal point, that is, the source of the earthquake, and travel in all directions through the Earth’s interior. If the Earth were an ideal elastic, homogeneous and isotropic body, the wavefronts of space waves would be concentric spheres centred at the epicentre of the earthquake. Spatial waves can be divided into so-called P-waves (from lat. primus = first) and S-waves (from lat. Secundus = second).
P-waves (primary waves) are longitudinal waves, which means that in them the particles oscillate around their equilibrium position in the direction of wave propagation (such as sound waves, for example). They are often also called compressional waves because they cause compression (compaction) of the means through which they pass. P-waves are the fastest type of seismic waves, with an approximate average speed of 6 km/s in the Earth’s crust and are therefore always the first to hit the receiver.
S-waves (secondary waves) are transverse waves and in them, the particles oscillate perpendicular to the direction of travel of the wave itself (for example, waves on water). They are also called shear waves because they cause the shearing of parts of the medium through which they spread. S-waves are slower, with an approximate speed of 3.5 km/s in the Earth’s crust, and they reach the receiver after P-waves, but their amplitudes can be many times higher than the amplitudes of P-waves, therefore they cause stronger shaking of the surface. Considering the geometric orientation of oscillations in space, i.e. polarization, SH- and SV-waves differ: SH-waves are S-waves polarized in the horizontal plane, while SV-waves are S-waves polarized in the vertical plane.
Unlike longitudinal waves that can propagate through a gas, liquid, or solid, transverse waves can only propagate through solids. This property is a key reason why we know that Earth’s outer core is liquid, not solid—S-waves cannot travel through it. S-waves are slower and reach the receiver only after P-waves. However, as their amplitudes are usually about five times higher than P-wave amplitudes, they will cause stronger shaking of the ground on the surface. Thus, during an earthquake, it is possible to feel two ground tremors: a weaker one caused by the arrival of P-waves and a stronger one caused by S-waves. Of course, if the focus of the earthquake is relatively close, both types of waves will arrive at the same time, so only one cherry will be felt. Likewise, with weaker earthquakes, most people will only feel the shaking caused by the S-waves.
Surface waves travel along the Earth’s surface and their amplitude decreases with depth. They are created by the interaction of space waves with discontinuities in the Earth’s interior. There are two types of surface waves: Love waves and Rayleigh waves.
Love waves (according to the British mathematician A.E.H. Love) cause the particles of the medium to move from one side to the other (animation). They are polarized in the horizontal plane, so they can only be seen on the horizontal components of the seismogram, and since they are faster, they reach the receiver before Rayleigh waves.
Rayleigh waves have both horizontal and vertical components of motion (according to the scientist John William Strutt better known as Lord Rayleigh). During the passage of the Rayleigh wave, the particles oscillate on ellipses; on the surface, they are ellipses whose horizontal axis is approximately equal to 2/3 of the vertical axis (animation). Oscillation can be retrograde (movements of particles on the “raised” part of the ellipse are in the same direction as the direction of wave propagation) or the opposite, directly.
Surface waves on the seismogram are best observed in the case of distant strong earthquakes whose focus is at depths less than 70 km. They are characterized by pronounced amplitudes, large periods, and fairly regular harmonic shapes and show the property of dispersion (their speeds depending on the frequency). Surface waves of periods between 15 and 30 s, which reach the seismological station last due to small groups of velocities, form the so-called Airy phase. On seismograms of distant earthquakes, this group very often has the largest amplitude and is used to define the magnitude of surface waves.
Seismic wave speeds are of the order of several kilometres per second (km/s) and since they depend on the type of material, based on seismograms it is possible to determine the composition of the part of the Earth’s interior through which the waves travel (speeds are lower if the material is denser and vice versa). Unfortunately, this process is not always that simple because sometimes different materials can have the same effect on the amount of velocity. Furthermore, the speed of seismic waves also depends on some other factors such as pressure and temperature. The pressure, which increases with depth due to the increasing mass in charge, causes the velocity to increase with depth. On the other hand, the increase in temperature towards the Earth’s interior will cause the speed of the waves to decrease. As the influence of pressure is more dominant than the influence of temperature, seismic velocities generally increase with increasing depth.
By the way, in the context of waves, two types of velocities are generally distinguished: phase velocity and group velocity. Phase velocity is the true velocity of wave propagation, while group velocity refers to the velocity of the maximum amplitude of a certain group of waves, i.e. the velocity with which its energy is propagated. The phenomenon that the phase and group velocity depend on the frequency of the wave is called dispersion. If the group speed is higher than the phase speed, it is a so-called anomalous dispersion. If the group velocity is lower than the phase velocity, the dispersion is normal. For the direct measurement of phase velocities, it is necessary to consider the arrival times of seismic waves at several stations. From the knowledge of the phase speed, it is possible to unambiguously determine the group speed, but the reverse is not valid. Group velocities can be determined if the time from the moment of the earthquake to the arrival of the wave at the station and the distance of the epicentre from the receiver are known. Thus, using the data of only one seismological station, it is relatively easy to determine the group velocity dispersion curve from the earthquake focus to the station. By comparing empirical values and theoretical group velocity curves, it is possible to define a model that describes the structure of the Earth’s interior.
A seismograph or seismometer is a measuring instrument used to measure and record ground motion. The surface of the Earth is constantly moving more or less under the influence of a whole series of factors such as earthquakes, explosions, volcanic activity, movement of air masses, human activity, the position of the Moon and other nearby bodies in the solar system, etc. The movement of the ground can be shown using physical quantities of displacement, speed, and acceleration. When measuring and recording ground motion, it is necessary to pay attention to the reference point against which the ground motion will be measured. Such a point must not be firmly connected to the ground because during an earthquake “everything” shakes. For this reason, a pendulum is used whose centre of gravity will remain stationary if the hanger experiences a sufficiently small displacement in a sufficiently short time. Unlike a pendulum, the seismograph housing moves freely, so it is possible to record the difference in their relative positions. The first seismographs worked on the principle of mechanical action, and today ground motion is most often converted into an electrical signal that is further processed and stored.
A seismogram is a record of a seismograph, i.e., a record of ground motion as a function of time. Seismographs usually simultaneously record three mutually perpendicular displacement components (vertical and two horizontal), so based on seismograms it is possible to completely reconstruct the ground motion of a certain area. Seismograms show the basic types of seismic waves generated by earthquakes and their numerous phases that reach the receiver after reflections and refractions at discontinuities in the Earth’s interior. The onset times of individual phases enable precise locating of the focus of the earthquake and insight into the structure of the part of the Earth through which the waves travelled.
In general, the recorded shifts on the seismogram are defined by the mechanism in the focus of the earthquake, the influence of the means through which the waves propagate, local shaping, and the action of the seismograph itself. Instrumental influence is the easiest to remove considering that it is defined by the construction of the measuring device. Local influences can also be removed relatively easily if the properties of local structures and assets near the station are known. To extract the contribution from the propagation of waves through the medium, simplified models are used that introduces a certain amount of unreliability. The reason for this is the fact that seismic waves when passing through the Earth’s interior, interact with the structures located in it – reflections, refractions, diffractions, conversions of one type of wave into another (conversion), increase or decrease in amplitude, etc. occur. All these effects are part of a complex displacement record on the seismogram, and it is very difficult to determine them individually, so modelling is a necessary step. By removing all the mentioned contributions, it is possible to single out the one related to the focal mechanism.
The focal mechanism describes the deformation field that generates seismic waves at the source of the earthquake. Most seismic sources involve faulting or shearing movements in the Earth’s interior, but some sources are caused by controlled underground explosions. The basic characteristic of sources caused by explosions is the spherical symmetry of the source itself and the displacement field, which is not observed in sources involving faulting. Faulting describes the relative displacement of two geological entities, the creation of a new fault or shearing along an already existing fault. The location, orientation, and type of fault can be determined using different methods (e.g., using the orientation of the first P-wave displacement recorded on the seismogram).
Seismic waves that arise in the focus of an earthquake are influenced by the distribution of deformations near the source itself, and they transmit this information to distant locations. Various mathematical models are used to represent the focal mechanism, such as the single or double-couple model, but also many other much more complex models. Furthermore, to geometrically describe the amplitudes of the initial motion in the vicinity of the source divided into P- and S-waves, the term wave radiation pattern is introduced. The location of the fault in space causes a certain distribution of radiation intensity. Therefore, predicting the relationship between the model of the spatial distribution of radiation, wave motion and the orientation of fault planes will enable the determination of remote fault processes that preceded the earthquake.
So, to determine the focal mechanism, seismologists represent the actual displacement field with an average displacement model that is simple enough to be replaced by a force system. In doing so, it is necessary to satisfy the condition that these forces are dynamically equivalent to the actual mechanism in the focal point, i.e., that they cause equivalent radiation of seismic waves.
Different parameters are used to represent the strength and impact of earthquakes on the Earth’s surface. One of them is the so-called macroseismic intensity. Macro seismic intensity describes the effect of the seismic source, which depends on the strength and distance of the earthquake and local soil properties. When calculating the intensity, the behaviour of objects on the surface and the way people experienced the event are initially taken into account. Using the distribution of macro seismic intensities, it is possible to determine the intensity at the source, i.e., the epicentre, and the depth of the focus of the earthquake. There are different scales for determining the macro seismic intensity, such as MCS (Mercalli-Cancani-Sieberg Scale), MSK (Medvedev-Sponheuer-Karnik Scale), MM (Modified Mercalli Scale), EMS (European Macroseismic Scale) and others.
In addition to intensity, magnitudes are used to estimate the power and released energy of an earthquake. Wanting to define a realistic magnitude that can be used in the assessment of seismic hazard, Charles F. Richter in 1935 defined the magnitude of an earthquake based on the magnitude used to describe the brightness of the stars. Magnitude is a measure used to describe the relative size/amount of released elastic energy of an earthquake. Its value derives from the largest displacement amplitude of the seismic wave recorded on the seismogram. The largest displacement amplitudes are precisely those that reflect the energy released in the focal point. It is important to note here that the energy released in the hypocentre of the earthquake is mostly spent on friction, i.e. that only a small part of it travels through space in the form of seismic waves. Furthermore, seismic waves will lose some of their energy along their path due to the increase of their wavefronts (geometric spreading) and properties of the medium (attenuation due to inelasticity and heterogeneity of the medium). Therefore, when defining the magnitude, it is necessary to carry out corrections related to energy losses.