CHAPTER: IV
EARTH QUAKES
Earthquakes: Causes of earthquakes; Effects of earthquakes; Types of Earthquake based on depth; Types of Earthquake/Seismic Hazards; Liquefaction; Tsunami; Intensity; Shield areas and seismic belts; Seismic wave; Seismic zones; Richter scale; Earthquake magnitude; Precautions to be taken for building construction in seismic areas; Frequency; Landslides; CREEP; Falls And Toppling; Slides; Flows; Earthflow; Causes of Landslides; Effects of Landslides; Aftershocks, foreshocks, and swarms; Importance of geophysical studies; Principles of geophysical analysis by Gravity Methods; Metamorphic Rocks; Electrical Methods; Resistivity Method; Electromagnetic Methods; Magnetic methods; Seismic Methods; Seismic-Refraction Method; Fundamental aspects of Rock Mechanics; Testing in rock mechanics; Environmental Geology; Importance of Environmental Geology; Importance of competence of sites by grouting; Microfine Cement; Permeable Reactive Barrier
Objectives:
Introduction
Seismology (the word is d6erived from the Greek word “seismos,” which means Earthquake, and logos mean science) is the scientific study of Earthquakes and the movement of waves through the Earth. The field also includes studies of variants such as seaquakes, volcanoes, and plate tectonics in general and consequential phenomena such as tsunamis. Engineering Seismology deals with the effects of Earthquakes on people and their environment and methods of reducing those effects. It is a very young discipline; many of its most significant developments have occurred in the past 30 to 40 years. Engineering Seismology requires knowledge of the geologic causes of and expected shaking, liquefaction, and other human effects, ranging from our buildings and other structures to the entire built and even social environment.
Earthquake, any sudden shaking of the ground caused by the passage of seismic waves through Earth’s rocks. Seismic waves are produced when some form of energy stored in Earth’s crust is suddenly released, usually when masses of rock straining against one another suddenly fracture and “slip.” Earthquakes occur most often along geologic faults, narrow zones where rock masses move about one another. The world’s major fault lines are located at the fringes of the enormous tectonic plates that make up Earth’s crust. (See the table of major earthquakes.)
Little was understood about earthquakes until the emergence of seismology at the beginning of the 20th century. Seismology, which involves the scientific study of all aspects of earthquakes, has yielded answers to such long-standing questions about why and how earthquakes occur. About 50,000 earthquakes large enough to be noticed without the aid of instruments occur annually over the entire Earth. Of these, approximately 100 are sufficient to produce substantial damage if their centers are near habitation areas. Very great earthquakes occur on average about once per year. Over the centuries, they have been responsible for millions of deaths and an incalculable amount of property damage.
Earth’s major earthquakes occur mainly in belts coinciding with the margins of tectonic plates. This has long been apparent from early catalogs of felt earthquakes and is even more readily discernible in modern seismicity maps, which show instrumentally determined epicenters. The most critical earthquake belt is the Circum-Pacific Belt, which affects many populated coastal regions around the Pacific Ocean—for example, New Zealand, New Guinea, Japan, the Aleutian Islands, Alaska, and the western coasts of North and South America. It is estimated that 80 percent of the energy presently released in earthquakes comes from those whose epicenters are in this Belt. However, the seismic activity is by no means uniform throughout the Belt, and there are several branches at various points. Because the Circum- Pacific Belt is associated with volcanic activity at many places, it has been popularly dubbed the “Pacific Ring of Fire.”
A second belt, known as the Alpide Belt, passes through the Mediterranean region eastward through Asia and joins the Circum-Pacific Belt in the East Indies. The energy released in earthquakes from this Belt is about 15 percent of the world’s total. There also are striking connected belts of seismic activity, mainly along oceanic ridges—including those in the Arctic Ocean, the Atlantic Ocean, and the western Indian Ocean—and along the rift valleys of East Africa. This global seismicity distribution is best understood in terms of its plate tectonic setting.
Earthquakes have varied effects, including changes in geologic features, damage to artificial structures, and impact on human and animal life. Most of these effects occur on solid ground, but since most earthquake foci are located under the ocean bottom, severe consequences are often observed along the margins of oceans.
Types of Earthquake based on depth
The most common type of Earthquake is the Tectonic Earthquake. These are produced when rocks break suddenly in response to various geological forces. Tectonic earthquakes are scientifically crucial to studying the Earth’s interior and tremendous social significance because they pose the most significant hazard.
95% of worldwide seismic energy release by plate tectonic and causes Tectonic earthquakes.
The point from which the seismic waves first emanate is called the earthquake focus or the hypocenter. The foci of natural earthquakes are at some depth below the ground surface. Shallow focus Earthquake – Those with foci less than 70 kilometers deep are called the external focus. Shallow earthquakes wreak the most devastation, and they contribute more than three-quarters of the total energy released in earthquakes throughout the world. Intermediate focus Earthquake – Those with foci from 70 to 300 kilometers deep are arbitrarily called the intermediate focus. Deep focus Earthquake – Those below the depth of 300 kilometers are termed deep focus. Many foci are situated hundreds of kilometers deep. Such regions include South American Andes, the Tonga islands, Somoa, The New Hebrides chain, the Japan Sea, Indonesia, and the Caribbean Antilles.
Earthquakes often cause dramatic geomorphologic changes, including ground movements— either vertical or horizontal—along geologic fault traces; rising, dropping, and tilting of the ground surface; changes in groundwater flow; liquefaction of sandy ground; landslides; and mudflows. The investigation of topographic changes is aided by geodetic measurements made systematically in several countries seriously affected by earthquakes.
Earthquakes can damage buildings, bridges, pipelines, railways, dams, and other structures. The type and extent of damage inflicted are related to the strength of the ground motions and the foundation soils’ behavior. In the most intensely damaged region, called the isoseismal area, the effects of a severe earthquake are usually complicated and depend on the topography and the nature of the surface materials. They are often more severe on soft alluvium and unconsolidated sediments than on hard rock. At distances of more than 100 km (60 miles) from the source, significant damage is caused by seismic waves traveling along the surface. There is frequently minor damage below depths of a few hundred meters in mines even though the ground surface immediately above is considerably affected.
Earthquakes are frequently associated with reports of distinctive sounds and lights. The sounds are generally low-pitched and have been likened to the noise of an underground train passing through a station. The occurrence of such sounds is consistent with the passage of high-frequency seismic waves through the ground. Occasionally, luminous flashes, streamers, and bright balls have been reported in the night sky during earthquakes. These lights have been attributed to electric induction in the air along with the earthquake source.
Types of Earthquake/Seismic Hazards
The effect of ground shaking – The first main earthquake hazard (danger) is the effect of ground shaking. Buildings can be damaged by shaking or the ground beneath them settling to a different level than before the Earthquake (subsidence). When an earthquake occurs, seismic waves will generate, and they will radiate away from the source and travels rapidly through the Earth’s crust. When these waves reach Earth’s surface, they will produce shaking, remaining for seconds to minutes. The strength and duration of shaking of a particular place depend upon the size and location of the Earthquake and the characteristics of the site. All the areas near the source of large Earthquakes will produce colossal damage. We can consider ground shaking as the most crucial hazard because all other risks occur due to ground shaking. Figure 1.1 shows the damage of buildings due to ground shaking.
Fig 1.1: Ground story collapse of a 4-storeyed building at Bhuj in Gujarat, India.
Ground displacement – The second main earthquake hazard is ground displacement (ground movement) along a fault. If a structure (a building, road, etc.) is built across a spot, the ground displacement during an earthquake could seriously damage or rip apart that structure. Faulting is the surface expression of the differential movement of blocks of the Earth’s crust. Faulting can be a simple “mole track” lateral move or a major vertical scrap or may not even be visible. Figure 1.2 shows typical ground displacement due to Earthquake.
Fig 1.2: Ground displacement
Flooding – An earthquake can rupture (break) dams or levees along a river. The water from the river or the reservoir would flood the area, damaging buildings and maybe sweeping away or drowning people. The first failure due to the Earthquake reported in the literature is Augusta dam, GA, during the 1886 Charleston, SC earthquake.
Fire – The fourth main earthquake hazard is fire. These fires can be started by broken gas lines and power lines or tipped over wood or coal stoves. They can be a severe problem, especially if the water lines that feed the fire hydrants are broken, too. For example, after the Great San Francisco Earthquake in 1906, the city burned for three days. Most of the town was destroyed, and 250,000 people were left homeless. Figure 3 shows the burning of San Francisco due to Earthquake.
Fig 1.3: San Francisco burning after the 1906 Earthquake
Landslide – Earthquakes can cause the failure of soil and rock slopes, particularly those that are marginally stable, to begin with. The most common type of landslides triggered by seismic events includes rock falls, soil slides, and rock slides on relatively steep slopes covered with disaggregated soil and rock. The debris from such failures can cut off roads and streams, damaging buildings, bridges, and other structures. Loss of life due to landslides is a common occurrence. Of particular concern are slope failures that progress quickly.
A tragic example is the Huascaran Mountain landslide that occurred as a result of the 1970 Peru earthquake. It buried the town of Yungay and part of the town of Ranrahirca, with a loss of life exceeding 18,000. Figure 1.4 shows typical landslide and slope failure due to the 2005 Pakistan earthquake.
Fig 1.4: Slope failure observed in the regions of Balakot and Muzaffarabad (Pakistan Earthquake October 8, 2005.)
Liquefaction is the temporary loss of soil strength and fluidization in certain saturated granular soils due to seismic shaking, representing a significant hazard in coastal areas and other locations with a high water table. The reduction in bearing capacity that accompanies the process of liquefaction can lead to the sinking of buildings, bridges, and other heavy structures, often with little or no damage to the system itself, as shown in Figures 1.5 and 1.6 below. Liquefied beds can sometimes be detected by the presence of surface sand boils, formed as mixtures of soil and water squeeze out of the ground, with the soil residue deposited nearby.
Fig 1.5: Ground failure due to liquefaction, Loma Prieta earthquake
Fig 1.6: Building Collapse due to liquefaction, 1964 Niigata earthquake, Japan
Tsunami – Tsunamis are long-period sea waves produced by rapid vertical seafloor movements caused by fault rupture during Earthquake. In the open sea, tsunamis travel great distances at high speed but are challenging to detect. They usually have a height less than 1 m and wavelength several hundreds of kilometers. As tsunami reaches onshore due to the water reduction, its speed will decrease, and size will increase. In some coastal areas, the seafloor’s shape is such that it will amplify the wave producing the vertical wall of water that rushes the outlying island and cause enormous damage. Figure 1.7 shows Tsunami hazards infamous Marian Beach, Chennai, India.
Figure 1.7: Destructions caused due to Tsunami in Marina Beach, Chennai
Some of the most common intensity scales are:
Magnitude
A quantitative measure is needed to compare the size of earthquakes worldwide, which is independent of the density of population and type of construction. Magnitude is the quantitative measurement of the amount of energy released by an earthquake.
Richter’s (1935) magnitude of a local earthquake is a logarithm base 10 of the maximum seismic wave amplitude recorded on a standard seismograph at a distance of 100 km from the epicenter. As earthquake sources are located at all distances from seismographic stations, Richter further developed a method of making allowance for the attenuation.
Richter scale magnitude is calibrated so that ML = 3 corresponds to an earthquake about a distance of 100 km with the maximum amplitude of A = 1mm. The most common modern magnitude scales are surface wave magnitude and body wave magnitude. However, Richter’s local volume does not distinguish between different types of waves (Table 1).
Table: 1. Understanding the Richter scale
Earthquakes are caused by the sudden release of energy within some limited region of the Earth’s rocks. The energy can be released by elastic strain, gravity, chemical reactions, or even the motion of massive bodies. Of all these, the release of elastic strain is the most critical cause because this form of energy is the only kind that can be stored in sufficient quantity on the Earth to produce significant disturbances. Earthquakes associated with this type of energy release are called tectonic earthquakes.
Tectonic earthquakes are explained by the so-called elastic rebound theory, formulated by the American geologist Harry Fielding Reid after the San Andreas Fault ruptured in 1906, generating the great San Francisco earthquake. According to the idea, a tectonic earthquake occurs when
strains in rock masses have accumulated to a point where the resulting stresses exceed the strength of the rocks and sudden fracturing results. The fractures propagate rapidly through the rock, usually tending in the same direction and sometimes extending many kilometers along a local zone of weakness. In 1906, for instance, the San Andreas Fault slipped along a plane 430 km (270 miles) long. Along this line, the ground was displaced horizontally as much as 6 meters (20 feet). As a fault rupture progresses along or up the fault, rock masses are flung in opposite directions and spring back to a position with less strain. At any one point, this movement may occur not at once but rather in irregular steps; these sudden slowings and restarting give rise to the vibrations that propagate as seismic waves. Such uneven properties of fault rupture are now included in the modeling of earthquake sources, both physically and mathematically. Roughness’s along the fault is referred to as asperities, and places, where the rupture slows or stops, are said to be fault barriers. Fault rupture starts at the earthquake focus, a spot that in many cases is close to 5–15 km under the surface. The crack propagates in one or both directions over the fault plane until stopped or slowed at a barrier. Sometimes, instead of being stopped at the fence, the fault rupture recommences on the far side; at other times, the stresses in the rocks break the border, and the crack continues. Earthquakes have different properties depending on the type of fault slip that causes them (as shown in the figure). The usual fault model has a “strike” (that is, the direction from north taken by a horizontal line in the fault plane) and a “dip” (the angle from the horizontal shown by the steepest slope in the fault). The lower wall of an inclined spot is called the footwall. Lying over the footwall is the hanging wall. When rock masses slip past each other parallel to the strike, the movement is known as strike-slip faulting. Movement parallel to the dip is called dip-slip faulting. Strike-slip faults are right-lateral or left lateral, depending on whether the block on the opposite side of the spot from an observer has moved to the right or left. In dip-slip defects, if the hanging-wall block moves downward relative to the footwall block, it is called “normal” faulting; the opposite motion, with the hanging wall moving upward close to the footwall, produces reverse or thrust faulting.
All known faults are assumed to have been the seat of one or more earthquakes in the past, though tectonic movements along faults are often slow, and most geologically ancient flaws are now aseismic (that is, they no longer cause earthquakes). The actual faulting associated with an earthquake may be complex, and it is often not clear whether in a particular earthquake the whole energy issues from a single fault plane. Observed geologic faults sometimes show relative displacements on the order of hundreds of kilometers over geologic time. In contrast, the sudden slip offsets that produce seismic waves may range from only several centimeters to tens of meters. In the 1976 Tangshan earthquake, for example, a surface strike-slip of about one meter was observed along the causative fault east of Beijing, and in the 1999 Taiwan earthquake, the Cheung-PU fault slipped up to eight meters vertically.
A separate type of Earthquake is associated with volcanic activity and is called a volcanic earthquake. Yet, it is likely that even in such cases, the disturbance is the result of a sudden slip of rock masses adjacent to the volcano and the consequent release of elastic strain energy. However, the stored energy may be of hydrodynamic origin due to heat provided by magma moving in reservoirs beneath the volcano or gas release under pressure.
There is a clear correspondence between the geographic distribution of volcanoes and significant earthquakes, particularly in the Circum-Pacific Belt and along oceanic ridges. Volcanic vents, however, are generally several hundred kilometers from the epicenters of most major shallow earthquakes, and many earthquake sources occur nowhere near active volcanoes. Even in cases where an earthquake’s focus occurs directly below structures marked by volcanic vents, there is
probably no immediate causal connection between the two activities; most likely, both result from the same tectonic processes.
Seismic zones in the Indian subcontinent are divided into four seismic zones (II, III, IV, and V) based on scientific inputs relating to seismicity, earthquakes that occurred in the past, and the tectonic setup region.
Previously, earthquake zones were divided into five zones concerning the severity of the earthquakes, but the Bureau of Indian Standards [IS 1893 (Part I):2002], has grouped the country into four seismic zones.; the first and second seismic zones were unified. The Bureau of Indian standards is the official agency for publishing the seismic hazard maps and codes. It has brought out seismic zoning maps: a six-zone map in 1962, a seven-zone map in 1966, and a five-zone map in 1970/1984.
Area with minor damage ( i.e., causing damages to structures with fundamentally periods more significant than 1.0 second ) earthquakes corresponding to intensities V to VI of MM scale (MM – Modified Mercalli Intensity scale). It covers the areas which are not covered by the other three seismic zones discussed below.
Moderate damage corresponding to intensity VII or MM scale. It comprises Kerala, Goa, Lakshadweep islands, Uttar Pradesh, Gujarat, and West Bengal, Parts of Punjab, Rajasthan, Madhya Pradesh, Bihar, Jharkhand, Chhattisgarh, Maharashtra, Orissa, Andhra Pradesh, Tamilnadu, and Karnataka.
Significant damage corresponding to intensity VII and higher of MM scale. It covers Jammu and Kashmir and Himachal Pradesh, National Capital Territory (NCT) of Delhi, Sikkim, Northern Parts of Uttar Pradesh, Bihar and West Bengal, parts of Gujarat, and small portions of Maharashtra near the west coast and Rajasthan.
Area determines by pro seismically of specific major fault systems. It is seismically the most active region and comprises entire northeastern India, parts of Jammu and Kashmir, Himachal Pradesh, Uttaranchal, and Rann of Kutch in Gujarat, North Bihar, and Andaman & Nicobar Islands. Earthquake zone V is the most vulnerable to earthquakes, where historically, some of the country’s most potent shocks have occurred. Earthquakes with magnitudes over 7.0 have occurred in these areas and have had intensities higher than IX (Fig:1.8).
Fig. 1.8: Seismic Zones in Indian Subcontinent and Intensity of the Map
The violence of seismic shaking varies considerably over a single affected area. Because the entire range of observed effects is not capable of simple quantitative definition, the strength of the vibration is commonly estimated by reference to intensity scales that describe the impact in qualitative terms. Intensity scales date from the late 19th and early 20th centuries, before seismographs capable of accurate measurement of ground motion, were developed. Since that time, the divisions in these scales have been associated with measurable accelerations of the local ground shaking. Intensity depends, however, is a complicated way not only on ground accelerations but also on the periods and other features of seismic waves, the distance of the measuring point from the source, and the local geologic structure. Furthermore, earthquake intensity, or strength, is distinct from earthquake magnitude, which measures the amplitude or size of seismic waves as specified by a seismograph reading. See below Earthquake magnitude.
Many different intensity scales have been set up during the past century and applied to current and ancient destructive earthquakes. For many years the most widely used was a 10- point scale devised in 1878 by Michele Stefano de Rossi and Franƈois-Alphonse Forel. The scale now generally employed in North America is the Mercalli scale, as modified by Harry O. Wood and Frank Neumann in 1931, in which intensity is considered to be more suitably graded. A 12- point abridged form of the modified Mercalli scale is provided below. Modified Mercalli intensity VIII is roughly correlated with peak accelerations of about one-quarter of gravity (g = 9.8 meters, or 32.2 feet, per second squared) and ground velocities of 20 cm (8 inches) second. Alternative scales have been developed in both Japan and Europe for local conditions. The European (MSK) scale of 12 grades is similar to the abridged version of the Mercalli.
Earthquake magnitude measures the “size,” or amplitude, of the seismic waves generated by an earthquake source and recorded by seismographs. (The types and nature of these waves are described in the section Seismic waves.) Because the size of earthquakes varies enormously, comparison purposes must compress the range of wave amplitudes measured on seismograms utilizing a mathematical device. In 1935 the American seismologist Charles F. Richter set up a magnitude scale of earthquakes as the logarithm to base 10 of the maximum seismic wave amplitude (in thousandths of a millimeter) recorded on a standard seismograph (the Wood-Anderson torsion pendulum seismograph) at a distance of 100 km (60 miles) from the earthquake epicenter. Reduction of amplitudes observed at various distances to the amplitudes expected at the standard length of 100 km is made based on empirical tables. Richter magnitudes ML are computed on the assumption that the ratio of the maximum wave amplitudes at two given distances is the same for all earthquakes and is independent of azimuth. Richter first applied his magnitude scale to shallow-focus earthquakes recorded within 600 km of the epicenter in the southern California region. Later, additional empirical tables were set up, whereby observations made at distant stations and on seismographs other than the standard type could be used. Finally, practical tables were extended to cover earthquakes of all significant focal depths and enable independent magnitude estimates from the body- and surface-wave observations. A current form of the Richter scale is shown in the table.
Table: 19. Richter scale for determination of earthquake magnitude
| Magnitude level |
Category |
Effects |
Earthquakes per year |
| less than 1.0 to 2.9 |
micro |
generally not felt by people, though recorded on local instruments |
more than 100,000 |
| 3.0–3.9 | minor | felt by many people; no damage | 12,000–100,000 |
| 4.0–4.9 | light | felt by all; minor breakage of objects | 2,000–12,000 |
| 5.0–5.9 | moderate | some damage to weak structures | 200–2,000 |
| 6.0–6.9 | strong | moderate damage in populated areas | 20–200 |
| 7.0–7.9 | major | serious damage over large areas; loss of life | 3–20 |
|
8.0 and higher |
great |
severe destruction and loss of life over large areas |
fewer than 3 |
At present, several different magnitude scales are used by scientists and engineers as a measure of the relative size of an earthquake. For one, the P-wave magnitude (Mb) is defined in terms of the amplitude of the P wave recorded on a standard seismograph. Similarly, the surface-wave extent (Ms) is determined in terms of the logarithm of the maximum amplitude of ground motion for surface waves with a wave period of 20 seconds.
At present, several different magnitude scales are used by scientists and engineers as a measure of the relative size of an earthquake. For one, the P-wave magnitude (Mb) is defined in terms of the amplitude of the P wave recorded on a standard seismograph. Similarly, the surface-wave extent (Ms) is determined in terms of the logarithm of the maximum amplitude of ground motion for surface waves with a wave period of 20 seconds.
Energy in an earthquake passing a particular surface site can be calculated directly from the recordings of seismic ground motion, given, for example, as ground velocity. Such recordings indicate an energy rate of 105 watts per square meter (9,300 watts per square foot) near a moderate-size earthquake source. The total power output of a rupturing fault in a shallow earthquake is on the order of 1014 watts, compared with the 105 watts generated in rocket motors. The surface-wave magnitude Ms has also been connected with the surface energy Es of an earthquake by empirical formulas. These give Es = 6.3 × 1011 and 1.4 × 1025 ergs for earthquakes of Ms = 0 and 8.9, respectively. A unit increase in Ms corresponds to approximately a 32-fold increase in energy. Negative magnitudes Ms correspond to the smallest instrumentally recorded earthquakes, a magnitude of 1.5 to the smallest felt earthquakes, and one of 3.0 to any shock felt at a distance of up to 20 km (12 miles). Earthquakes of magnitude 5.0 cause light damage near
the epicenter. Those of 6.0 are destructive over a restricted area, and those of 7.5 is at the lower limit of major earthquakes. The total annual energy released in all earthquakes is about 1025 ergs, corresponding to a rate of work between 10 million and 100 million kilowatts. This is approximately one one-thousandth the annual amount of heat escaping from the Earth’s interior. Ninety percent of the total seismic energy comes from earthquakes of magnitude 7.0 and higher—that is, those whose energy is on the order of 1023 ergs or more.
There also are empirical relations for the frequencies of earthquakes of various magnitudes. Suppose N to be the average number of shocks per year for which the importance lies in a range about Ms. Thenlog10 N = a − benefits the data well both globally and for particular regions; for example, for shallow earthquakes worldwide, a = 6.7 and b = 0.9 when Ms > 6.0. Therefore, the frequency for more significant earthquakes increases by a factor of about 10 when the magnitude is diminished by one unit. However, the increase in frequency with the reduction in Ms falls short, however, of matching the decrease in the energy E. Thus, more significant earthquakes are overwhelmingly responsible for most of the total seismic energy release. The number of earthquakes per year with Mb > 4.0 reaches 50,000.
Earth flows range from very small to very big, involving hundreds of tons of material blocking or destroying roads, damming rivers, and destroying houses.
Debris Flow or Mudflow
These two terms are used interchangeably, and they refer to the rapid but viscous flow of mud and other surficial materials. Rotational slides usually end up as mudflows after traveling a few meters because the soil is saturated. The vibrations caused by the movement induce the Earth to liquefy and behave as a viscous fluid. The flow can travel along channels or flow paths for considerable distances until the slope decreases or the channel widens, pointing the flow fans out and its momentum decreases. Mud or debris flows commonly originate in steep terrain where vegetation and organic litter that helps to stabilize the soil and retain rainfall and runoff have been removed by fire, grazing, logging, or other processes. Intense and prolonged rainfall may then trigger the downslope movement of soil and other surface materials. This type of landslide is potentially more dangerous than different types because it can form quickly and more velocities up to 80km per hour. The greater density meant that it is more destructive than floodwaters. The mud does not recede after the storm. In the unlikely event of a volcanic eruption in Fiji, mudflows can be generated on its flanks by rapid infusion of significant volumes of water (from heavy rains associated with such outbreaks) into poorly consolidated ash and other volcanic debris deposits (Fig. 4.6).
Figure: 4.6 Debris flow
CREEP
Creep occurs mainly in the soil mantle, that part of the soil from the surface to a few centimeters or meters below the surface. It involves the slow downslope movement or the gradual plastic deformation of the soil mantle and the fracturing of bedrock at imperceptible rates (Fig 4.7). There is no single surface along which slippage occurs. The downhill movement or creep rate can vary from a few millimeters per year for slopes less than 10% to about 10mm per year in steeper terrains. The downward movement involves the minute displacement of individual particles that are moving at different rates. It is commonly caused by the expansion of the surface layer due to heating followed by contraction due to cooling. Creep may also be caused by the swelling of certain clays after seasonal rainfalls when their moisture content increases, followed by contraction when their moisture content drops during the dry period.
Figure: 4.7 Creep
Falls And Toppling A rockfall is the abrupt free fall or downslope movement (rolling or sliding) of loosened blocks or boulders of solid rock. It differs from a slide in that free fall is the primary type of movement, and no marked slide surface develops. This type of slope failure occurs in caverns, steep gorges, sea cliffs, and steep road cuts through unstable bedrock. The bedding, jointing, and fracturing of the bedrock are the critical factors affecting slope stability. In addition, the effects of weathering, such as the freezing of water in joints (in cold countries), the pressure of water in fissures, and root pressures, may initiate failure in the weak rocks. A rockfall, as in most landslides, is usually the result of a combination of factors. For example, on a sea cliff, it could be due to various jointing patterns, percolation of surface water, wedging of tree roots, and the impact of and undercutting by waves. Thus a lot of rock falls along sea cliffs occur during storms when much rain percolates through cracks in the rock, and the pressure pushes the blocks over or when heavy surf strikes the ridge causing vibrations and thus causing undercut cliffs faces to topple over. The magnitude and scale of rockfalls vary from the breaking off small isolated rocks to the fall of enormous masses. Largescale failures have been known to dam rivers, creating lakes, and destroying parts of towns. On a small scale, the talus commonly found at the base of cliffs and the bottom of slopes in mountainous areas is accumulating numerous rockfalls over many years (Fig.4.8).
Figure: 4.8 Rockfall and Topple landslides
Slides A slide is characterized by failure of material at depth and then moved by sliding along a rupture or slip surface in the strictest sense. If sliding is on a predominantly planar slip surface, then the slide is called a block slide. If movement is on a curved slip surface, then the slide is called a rotational slide. Many rotational slides end up as mudflow leaving a gaping hole in the ground where the slide began. Debris from the slide is strewn down a torrent track along which the mudflow traveled to the base of the slope or where the flow path widens and dissipates. A rotational slide with one or more curved slip surfaces where the movement of material is incomplete, leaving individual slumped blocks, is referred to as a slump (Fig: 4.9).
Figure: 4.9 Translational, Rotational, block, and Lateral landslides
Slides are probably the most common and, overall, possibly the most destructive type of landslide to hillside developments. Wherever steep mountains or hillside slopes occur or are altered, the possibility of large landslides and consequent disasters exist. The rupture or slip surface can appear within the bedrock, at the contact between the bedrock and the overburden or soil (in which case all the surface materials move), or within the overburden, which in some cases may be of artificial fill.
Flows
Flows involve the deformation of an entire soil mass that then flows downslope as a viscous or sticky fluid. Deformation may be due to a high soil water content or seismic shaking, leading to liquefaction and generating such a fluid flow. The slopes need not be very steep. Two types of flow can be recognized; if the downslope movement is prolonged, then it is an earthflow; if it is very rapid, it is a debris flow or, as it is sometimes known, a mud flow.
Earthflow
Earth flows occur in moderate to steep slopes where the topsoil or overburden seasonally becomes saturated by heavy rains. As a result, the material slumps away from the upper part of the slope, leaving a scarp, and flows down to form a bulge at the toe (Fig: 4.10).
Figure: 4.10 Earthflow landslide
Faults may pose a severe problem if open to the passage of water. They become potential outlets for the escape of stored water from the reservoir. They can be treated by grouting or trenching along the fracture line and filling the trench with clay puddle or concrete. Landslides are indications of an unstable state. Such grounds known to have been subjected to landslides should be avoided. Water leaking through a porous bed may lead to landslides on slopes away from the reservoir sometime after the pool is filled. Global seismicity patterns had no robust theoretical explanation until the dynamic model called plate tectonics was developed during the late 1960s. This theory holds that the Earth’s upper shell, or lithosphere, consists of nearly a dozen large, quasi-stable slabs called plates. The thickness of each of these plates is roughly 80 km (50 miles). The plates move horizontally relative to neighboring plates at a rate of 1 to 10 cm (0.4 to 4 inches) per year over a shell of lesser strength called the asthenosphere. At the plate edges where there is contact between adjoining plates, boundary tectonic forces operate on the rocks, causing physical and chemical changes in them. The new lithosphere is created at oceanic ridges by the upwelling and cooling of magma from the Earth’s mantle. The horizontally moving plates are believed to be absorbed in the ocean trenches, where a subduction process carries the lithosphere downward into the Earth’s interior. The total amount of lithospheric material destroyed at these subduction zones equals that generated at the ridges.
Seismological evidence (such as the location of significant earthquake belts) is everywhere in agreement with this tectonic model. Earthquake sources are concentrated along the oceanic ridges, which correspond to divergent plate boundaries. At the subduction zones, which are associated with convergent plate boundaries, intermediate- and deep-focus earthquakes mark the location of the upper part of a dipping lithosphere slab. The focal mechanisms indicate that the stresses are aligned with the dip of the lithosphere underneath the adjacent continent or island arc. Some earthquakes associated with oceanic ridges are confined to strike-slip faults, called transform faults, that offset the ridge crests. Slip motions characterize the majority of the earthquakes occurring along such horizontal shear faults. Also in agreement with the plate tectonics theory is the high seismicity encountered along the edges of plates where they slide past each other. Plate boundaries of this kind, sometimes called fracture zones, including the San Andreas Fault in California and the North Anatolian fault system in Turkey. Such plate boundaries are the site of interpolating earthquakes of shallow focus.
The low seismicity within plates is consistent with the plate tectonic description. Small to large earthquakes occur in limited regions well within the boundaries of plates; however, such interpolate seismic events can be explained by tectonic mechanisms other than plate boundary motions and their associated phenomena.
Earth’s major earthquakes occur mainly in belts coinciding with the margins of tectonic plates. This has long been apparent from early catalogs of felt earthquakes and is even more readily discernible in modern seismicity maps, which show instrumentally determined epicenters. The most crucial earthquake belt is the Circum-Pacific Belt, which affects many populated coastal regions around the Pacific Ocean—for example, New Zealand, New Guinea, Japan, the Aleutian Islands, Alaska, and the western coasts of North and South America. It is estimated that 80 percent of the energy presently released in earthquakes comes from those whose epicenters are in this Belt. However, the seismic activity is by no means uniform throughout the Belt, and there are several branches at various points. Because the Circum- Pacific Belt is associated with volcanic activity at many places, it has been popularly dubbed the “Pacific Ring of Fire.”
A second belt, known as the aliped Belt, passes through the Mediterranean region eastward through Asia and joins the Circum-Pacific Belt in the East Indies. The energy released in earthquakes from this Belt is about 15 percent of the world’s total. There also are striking connected belts of seismic activity, mainly along oceanic ridges—including those in the Arctic Ocean, the Atlantic Ocean, and the western Indian Ocean—and along the rift valleys of East Africa. This global seismicity distribution is best understood in terms of its plate tectonic setting.
Earthquakes have varied effects, including changes in geologic features, damage to artificial structures, and impact on human and animal life. Most of these effects occur on solid ground, but since most earthquake foci are located under the ocean bottom, severe consequences are often observed along the margins of oceans.
Earthquakes often cause dramatic geomorphologic changes, including ground movements— either vertical or horizontal—along geologic fault traces; rising, dropping, and tilting of the ground surface; changes in the flow of groundwater; liquefaction of sandy ground; landslides; and mudflows. The investigation of topographic changes is aided by geodetic measurements made systematically in several countries seriously affected by earthquakes.
Earthquakes can damage buildings, bridges, pipelines, railways, dams, and other structures. The type and extent of damage inflicted are related to the strength of the ground motions and the behavior of the foundation soils. In the most intensely damaged region, called the Meizu seismal area, the effects of a severe earthquake are usually complicated and depend on the topography and the nature of the surface materials. They are often more severe on soft alluvium and unconsolidated sediments than on hard rock. At distances of more than 100 km (60 miles) from the source, the primary damage is caused by seismic waves traveling along the surface. There is frequently minor damage below depths of a few hundred meters in mines even though the ground surface immediately above is considerably affected.
Earthquakes are frequently associated with reports of distinctive sounds and lights. The sounds are generally low-pitched and have been likened to the noise of an underground train passing through a station. The occurrence of such sounds is consistent with the passage of high-frequency seismic waves through the ground. Occasionally, luminous flashes, streamers, and bright balls have been reported in the night sky during earthquakes. These lights have been attributed to electric induction in the air along with the earthquake source.
Most parts of the world experience occasional shallow landslides, originating within 60 km (40 miles) of the Earth’s outer surface. Thus, the great majority of earthquake foci are shallow. However, it should be noted that the geographic distribution of more minor earthquakes is less completely determined than more severe quakes, partly because the availability of relevant data depends on the distribution of observatories.
Of the total energy released in earthquakes, 12 percent comes from intermediate earthquakes, quakes with a focal depth ranging from about 60 to 300 km. About 3 percent of total energy comes from deeper earthquakes. The frequency of occurrence falls off rapidly with increasing focal depth in the intermediate range. Below intermediate-depth, the distribution is relatively uniform until the greatest focal depths, of about 700 km (430 miles), are approached.
The deeper-focus earthquakes commonly occur in patterns called Benioff zones that dip into the Earth, indicating the presence of a subducting slab. Dip angles of these slabs average about 45°, with some shallower and others nearly vertical. Benioff zones coincide with tectonically active island arcs such as Japan, Vanuatu, Tonga, and the Aleutians. They usually are but not always associated with deep ocean trenches such as those along the South American Andes. Exceptions to this rule include Romania and the Hindu Kush mountain system. In most Benioff zones, intermediate- and deep-earthquake foci lie in a thin layer. However, recent precise hypocentral locations in Japan and elsewhere show two distinct parallel bands of fetishes 20 km apart.
Usually, a primary or even moderate earthquake of shallow focus is followed by many lesser-size earthquakes close to the source region. This is to be expected if the fault rupture producing a significant earthquake does not relieve all the accumulated strain energy at once. This dislocation is liable to cause an increase in the stress and strain at several places in the vicinity of the focal region, bringing crustal rocks at specific points close to the pressure at which fracture occurs. In some cases, an earthquake may be followed by 1,000 or more aftershocks a day.
Sometimes a large earthquake is followed by a similar fault source within an hour or perhaps a day. An extreme case of this is multiple earthquakes. However, in most instances, the first principal Earthquake of a series is much more severe than the aftershocks. In general, the number of aftershocks per day decreases with time. The aftershock frequency is roughly inversely proportional to the time since the largest Earthquake of the series.
Most significant earthquakes occur without detectable warning, but foreshocks precede some principal earthquakes. In another typical pattern, large numbers of small earthquakes may occur in a region for months without a major earthquake. For instance, in the Matsushiro region of Japan, there occurred between August 1965 and August 1967 a series of hundreds of thousands of earthquakes, some sufficiently strong (up to Richter magnitude 5) to cause property damage but no casualties. The maximum frequency was 6,780 small earthquakes on April 17, 1966. Such a series of earthquakes are called earthquake swarms. Earthquakes associated with volcanic activity often occur in swarms, though swarms also have been observed in many nonvolcanic regions.
Seismic waves generated by an earthquake source are commonly classified into three main types. The first two, the P (or primary) and S (or secondary) waves, propagate within the body of the Earth, while the third, consisting of Love and Rayleigh waves, propagate along its surface. These seismic waves were mathematically predicted during the 19th century, and modern comparisons show a close correspondence between such theoretical calculations and actual measurements of the seismic waves. The P seismic waves travel as elastic motions at the highest speeds. Thus, longitudinal waves can be transmitted by both solid and liquid materials in the Earth’s interior. With P waves, the particles of the medium vibrate like sound waves—the transmitting media is alternately compressed and expanded. The slower type of body wave, the S wave, travels only through solid material. With S waves, the particle motion is transverse to the direction of travel and involves shearing off the transmitting rock. Because of their more incredible speed, P waves are the first to reach any point on the Earth’s surface. The first P-wave onset starts from the spot where an earthquake originates. This point, usually at some depth within the Earth, is called the focus or hypocentre. The point at the surface immediately above the focus is known as the epicenter. The free surface of the Earth guides loves and Rayleigh waves. They follow along after the P and S waves have passed through the body of the planet. Both Love and Rayleigh waves involve horizontal particle motion, but only the latter type has vertical ground displacements. As Love and Rayleigh’s waves travel, they disperse into long wave trains, and, at substantial distances from the source in alluvial basins, they cause much of the shaking felt during earthquakes.
The gravity field of the Earth can be measured by timing the free fall of an object in a vacuum, measuring the period of a pendulum, or various other ways. Today almost all gravity surveying is done with gravimeters. Such an instrument typically consists of a weight attached to a spring that stretches or contracts, corresponding to an increase or decrease in gravity. It is designed to measure differences in gravity accelerations rather than absolute magnitudes. Gravimeters used in geophysical surveys have an accuracy of about 0.01 milligram (mg; 1 mg
= 0.001 centimetre per second per second). That is to say, they are capable of detecting differences in the Earth’s gravitational field as small as one part in 100,000,000.
Gravity differences over the Earth’s surface occur because of local density differences between adjacent rocks. The variations in the density of the crust and cover are presented on a gravity anomaly map. A gravity anomaly map looks at the difference between the value of gravity measured at a particular place and the predicted value for that place. Gravity anomalies form a pattern, which may be mapped as an image or by contours. The wavelength and amplitude of the gravity anomalies give geoscientists an idea of the size and depth of the geological structures causing these anomalies. Deposits of very dense and heavy minerals will also affect gravity at a given point and produce an anomaly above normal background levels.
Anomalies of exploration interest are often about 0.2 mg. Data must be corrected for variations due to elevation (one meter is equivalent to about 0.2 mg), latitude (100 meters are equal to about 0.08 mg), and other factors. Gravity surveys on land often involve meter readings every kilometer along with traverse loops a few kilometers across. It takes only a few minutes to read a gravimeter, but determining location and elevation accurately requires much effort.
Gravity measurements can be obtained either from airborne (remote) or ground surveys. However, the most sensitive surveys are currently achieved from the ground. This is because variations of gravity are due to local changes in rock density and depend on the type of rocks beneath the surface. Sedimentary rocks are, for example, less dense than GraniteGranite, which is, in turn, less dense than basalt.
High Density
Extrusive Igneous Rocks, Eg. Basalt
For the human exploration of the solar system, instruments must meet the criteria of low mass, low volume, low power demand, safe operation, and ruggedness and reliability (Meyer et al., 1995; Hoffman, 1997; Budden, 1999). In addition, tools used for planetary exploration will need to address fundamental scientific questions and identify precious resources, such as water.
The primary goal of studying detailed gravity data is to provide a better understanding of subsurface geology. The gravity method is a relatively cheap, non-invasive, non-destructive remote sensing method that has already been tested on the lunar surface. It is also passive – that is, no energy needs to be put into the ground to acquire data; thus, the method is well suited to a populated setting such as Taos and a remote location such as Mars. The small portable instrument also permits walking traverses – ideal because of the congested tourist traffic in Taos.
Measurements of gravity provide information about the densities of rocks underground. There is a wide range in thickness among rock types, and therefore geologists can make inferences about the distribution of strata. For example, in the Taos Valley, we are attempting to map subsurface faults. Because faults commonly juxtapose rocks of differing densities, the gravity method is an excellent exploration choice (Fig: 4.5).
Gravity Survey – Measurements of the gravitational field at a series of different locations over an area of interest. The objective in exploration work is to associate variations with differences in the distribution of densities and hence rock types.
| Figure: 4.5 Cartoon illustrations showing the relative surface variation of gravitational acceleration over geologic structures. |
Metamorphic Rocks
Intrusive Igneous Rocks, e.g., GraniteGranite
In most cases, the density of sedimentary rocks increases with depth because increasing pressure reduces porosity. Uplifts usually bring denser rocks nearer the surface and thereby create positive gravity anomalies. Faults that displace rocks of different densities also can cause gravity anomalies. Salt domes generally produce negative anomalies because salt is less dense than the surrounding rocks. Such faults, folds, and salt domes trap oil, and so the detection of gravity anomalies associated with them are crucial in petroleum exploration. Moreover, gravity measurements are occasionally used to evaluate the amount of high-density minerals present in an orebody. They also provide a means of locating hidden caverns, old mine workings, and other underground cavities.
Several other factors also control density contrasts of different materials. The most important is the grain density of the particles forming the material, the porosity of the material, and the interstitial fluids within the fabric—generally, specific gravities of soil and shale range from 1.7 to 2.2. Massive limestone averages 2.7. While this range of values may appear to be pretty significant, local contrasts will be only a fraction of this range. A typical order of magnitude for regional density contrasts is 0.25.
Gravity surveys provide an inexpensive method of determining regional structures associated with groundwater aquifers or petroleum traps. Gravity surveys have been one of the principal exploration tools in regional petroleum exploration surveys. However, gravity surveys have somewhat limited applications in geotechnical investigations.
Electrical Methods:
Electrical methods are used to map variations in electrical properties of the subsurface. The main physical property involved is electrical conductivity, which measures how easily an electrical current can pass through a material. Subsurface materials exhibit an extensive range of electrical conductivity values. New rock is generally a poor conductor of electricity, but a select group of metallic minerals containing iron, copper, or nickel is an excellent conductor. Layers of graphite are also excellent conductors.
The examples of good conductors mentioned above are pretty rare. For most rocks, the electrical conductivity is mainly governed by the amount of water filling the pore spaces and the amount of salt dissolved in this water. Pure water has a very low electrical conductivity. On the other hand, seawater, which contains high levels of dissolved salts such as NaCl, is a relatively good conductor of electrical current. Groundwater can vary in salt content from fresh through brackish (slightly salty) to saline (similar in salt content to seawater) through to hyper-saline (saltier than seawater).
The electrical conductivity of rocks is not the only attribute that is of value to exploration geologists. Several different electrical properties of stones are measured and interpreted in mineral exploration. They depend on:
Self Potential Method: Some materials tend to become natural batteries that generate natural electric currents whose effects can be measured. The self-potential method relies on the oxidation of the upper surface of metallic sulfide minerals by downward-percolating groundwater to become a natural battery; current flows through the ore body and back through the surrounding groundwater, which acts as the electrolyte. Measuring the biological voltage differences – usually 50-400 millivolts (mV), permits detecting metallic sulfide bodies that lie above the water table. Other mineral deposits that can generate self-potentials are graphite, magnetite, anthracite, and pyritized rocks. Induced Polarization: The passage of an electric current across an interface where conduction changes from ionic to electronic results in a charge buildup at the interface. This charge builds up shortly after the current flow begins, and it takes a short time to decay after the current circuit is broken. Such an effect is measured in induced-polarization methods and is used to detect sulfide ore bodies.
Resistivity Method
Resistivity methods involve passing a current from a generator or other electric power source between a pair of current electrodes and measuring potential differences with another pair of electrodes. Various electrode configurations are used to determine the apparent resistivity from the voltage/current ratio. The resistivity of most rocks varies with porosity, the salinity of the interstitial fluid, and certain other factors. For example, rocks containing appreciable clay usually have low resistivity. The resistivity of rocks containing conducting minerals such as sulfide ores and graphitized or pyritized rocks depends on the connectivity of the minerals present. Resistivity methods are also used in engineering and groundwater surveys because resistivity often changes markedly at soil/bedrock interfaces, water tables, and a fresh/saline water boundary.
Electrical Resistivity Method:
The electrical resistivity method consists of measuring the resistivity of the soil strata and correlating the resistivity to the properties of the soil. The principal application of the electrical resistivity method is in investigating foundations of dams and other large structures, particularly in exploring granular river channel deposits or bedrock surfaces. The technique is also used for locating fresh or saltwater boundaries.
The electrical resistivity method is of the following two types:
Electrical Profiling Method:
In this method, four electrodes, usually in metal spikes, are driven into the ground at the same spacing. The two outer electrodes are known as current electrodes, and the two inner electrodes are known as potential electrodes, as shown in Fig. 14.22.
A direct current (DC) of 50-100 milliamperes (mA) is applied between the outer electrodes, and the voltage drop or the potential difference between the inner electrodes is measured using a potentiometer. The mean resistivity of the soil up to a depth of D cm below ground surface is obtained from Eq. (14.16) as follows –
Where D is the distance between electrodes in centimeters (cm), V the voltage drop between inner electrodes in volts (V), and I the current flowing between outer electrodes in amperes (A).
The electrodes are moved as a group, at the same spacing between them, as shown in Fig. 14.23(a), and different profile lines are run across the area. The test is repeated after changing the spacing to determine the mean resistivity up to a depth equal to the new spacing. The method is also known as the resistivity mapping method. The type of soil or rock stratum encountered will be estimated using the measured resistivity from Table 14.9.
This method is similar to the electrical profiling method, except that the electrode system is expanded about a point P by increasing the spacing between the electrodes in successive operations. Thus, the electrode spacing increases with every subsequent test, as shown in Fig. 14.23(b).
As the depth of current penetration is equal to the electrode spacing, the change in the mean resistivity is correlated to the changes in the strata at that location. As a result, the midpoint (P) moves forward by a distance of 4D in successive tests of the electrical profiling method. However, the middle (P) remains at the same position in the electrical sounding process.
Limitations:
The electrical resistivity method is less reliable than the seismic-refraction method since the resistivity of a particular soil or rock can vary over a wide range of values depending on the density, voids or fractures, and degree of saturation soil.
Limitations of the electrical resistivity method are the following:
iii. As the resistivity of different strata at the interface changes gradually and not abruptly, the interpretation becomes difficult.
Electromagnetic Methods
The passage of current in the general frequency range of 500-5,000 hertz (Hz) induces the Earth electromagnetic waves of long wavelength, which have considerable penetration into the Earth’s interior. The effective penetration can be changed by altering the frequency. Eddy currents are induced where conductors are present. These currents generate an alternating magnetic field, which causes a secondary voltage that is out of phase with the primary voltage in a receiving coil. Electromagnetic methods involve measuring this out-of-phase component or other effects, making it possible to locate low-resistivity ore bodies wherein the eddy currents are generated.
Several electrical methods described above are used in boreholes. The self-potential (SP) log indicates mainly clay (shale) content because an electrochemical cell is established at the shale boundary when the salinity of the borehole (drilling) fluid differs from that of the water in the rock. Resistivity measurements are made by using several electrode configurations and also by induction. Borehole methods are used to identify the rocks penetrated by a borehole and determine their properties, especially their porosity and the nature of their interstitial fluids.
Magnetic methods
One of the essential tools in modern mineral exploration methods is the magnetic survey. Magnetic surveys are fast, provide a great deal of information for the cost, and provide information about the distribution of rocks occurring under thin layers of sedimentary rocks – useful when trying to locate orebodies. When the Earth’s magnetic field interacts with a magnetic mineral contained in a rock, the stone becomes magnetic. This is called induced magnetism. However, a stone may be magnetic if at least one of its minerals is composed of is magnetic. The strength of a rock’s magnetism is related to the number of magnetic minerals it contains and the physical properties, such as grain size, of those minerals. The main magnetic mineral is magnetite (Fe3O4), which is disseminated through most rocks in differing concentrations. Measurements of the Earth’s total magnetic field or any of its various components can be made. The oldest magnetic prospecting instrument is the magnetic compass, which measures the field direction. Other devices, which are appreciably more accurate, include magnetic balances, fluxgate magnetometers, proton-precession, and optical-pumping magnetometers. Magnetic effects result primarily from the magnetization induced in susceptible rocks by the Earth’s magnetic field. Most sedimentary rocks have shallow susceptibility and thus are nearly transparent to magnetism. Accordingly, in petroleum exploration, magnetic surveys are used negatively – magnetic anomalies indicate the absence of explorable sedimentary rocks. Magnetic surveys map features in igneous and metamorphic rocks, possible faults, dikes, or other features associated with mineral concentrations. Data are usually displayed in the form of a contour map of the magnetic field, but the interpretation is often made on profiles. It must be remembered that rocks cannot retain magnetism when the temperature is above the Curie point (» 500oC for most magnetic materials), and this restricts magnetic stones to the upper 40 kilometers of the Earth’s interior.
When exploring petroleum, magnetic surveys are usually done with magnetometers borne by aircraft flying in parallel lines spaced two to four kilometers apart at an elevation of about 500 meters. When searching for mineral deposits, the flight lines are spaced 0.5 to 1.0 kilometers apart at the height of roughly 200 meters above the ground. Ground surveys are conducted to follow up magnetic anomalies identified through aerial surveys. Such surveys may involve stations spaced only 50 meters apart. A ground monitor is usually used to measure the natural fluctuations of the Earth’s field overtime to make corrections. Surveying is generally suspended during periods of sizeable magnetic flux (magnetic storms).
Seismic Methods
Seismic methods are based on measurements of the time interval between the initiation of a seismic (elastic) wave and its arrival at detectors. The seismic wave may be generated by an explosion, a dropped weight, a mechanical vibrator, a bubble of high-pressure air injected into water, or other sources. A Geophone detects the seismic wave on land or by a hydrophone in water. An electromagnetic Geophone generates a voltage when a seismic wave produces relative motion of a wire coil in the field of a magnet. In contrast, a ceramic hydrophone generates a voltage when deformed by the passage of a seismic wave. Data are usually recorded on magnetic tape for subsequent processing and display. Seismic methods are of two kinds – Refraction methods and Reflection methods.
Seismic refraction methods
Seismic energy travels from source to detector by many paths. When near the start, the initial seismic energy generally travels by the shortest route, but as the source to geophone distances become greater, seismic waves traveling by longer routes through rocks of higher seismic velocity may arrive earlier. Such waves are called head waves, and the refraction method involves their interpretation. From a plot of travel time as a function of the source to geophone distance, the number, thicknesses, and velocities of rock layers present can be determined for simple situations. The assumptions usually made are that:
The velocity values determined from time-distance plots also depend on the dip (slope) of interfaces, apparent velocities increasing when the geophones are up-dip from the source and decreasing when down drop. By measuring in both directions the depth and rock velocity, each can be determined. With sufficient measurements, relief on the interfaces separating the layers also can be ascertained.
High-velocity bodies of local extent can be located by fan shooting. Travel times are measured along different azimuths from a source, and an abnormally early arrival time indicates that a high-velocity body was encountered at that azimuth. This method has been used to detect salt domes, reefs, and intrusive bodies characterized by higher seismic velocity than the surrounding rock. In addition, seismic waves may be used for various other purposes. They are employed, for example, to detect faults that may disrupt a coal seam or fractures that may allow water penetration into a tunnel.
Seismic-Refraction Method:
The seismic-refraction method is based on the principle that elastic shock waves travel at different velocities in different materials. Shock waves are generated at a point on the ground surface using a sled hammer. These waves travel deep into the ground and get refracted at the interface of two different materials and to the ground surface. The arrival of these waves at various locations on the ground surface is recorded by geophones, which pick up the refracted waves. The geophones convert the ground vibrations into electrical impulses and transmit them to a recording apparatus.
When the distance between the vibration source and the geophone is short, the arrival time will be a direct wave. However, when the length exceeds a specific value (depending on the thickness of the stratum), the refracted wave will be the first to be detected by the geophone [see Fig. 14.21(a)]. This is because the refracted wave, although more extended than the direct wave, passes through a stratum of higher density (and hence higher seismic velocity).
A plot is made between distance on the x-axis and time on the y-axis, as shown in [Fig. 14.21(b)]. Points B and C in Fig. 14.21(b) represent the distance at which the refracted wave from the second and third strata arrive at the geophone, marked by a change in the slope of the graph. The slope of line AB gives the reciprocal of seismic velocity in the top layer (1/v1, that of BC gives (1/v2), and that of CD gives (1/v3), etc. The thickness (H1) of the top stratum of the soil is given by
D1 is the distance corresponding to point B, where the seismic velocity changes from v1 to v2.
If there is a slight or considerable variation in the thickness of the top stratum, H1 represents the average thickness. Shepard and Haines (1944) gave Eqs. (14.12) – (14.15) for determining depths H1 and H2 in a three-layer stratum.
where l1 is the length AB1 = Intercept of line AB on the y-axis and l2 = AC1 – AB1 = B1C1 [see Fig. 14.21(b)].
The seismic-refraction method is based on the following assumptions:
iii. The boundaries between strata are distinct horizontal or inclined planes.
Seismic reflection methods
Most seismic work utilizes reflection techniques. Sources and geophones are essentially the same as those used in refraction methods. The concept is similar to echo sounding – seismic waves are reflected at interfaces where rock properties change. The round-trip travel time, together with velocity information, gives the distance to the interface. The relief on the interface can be determined by mapping the reflection at many locations. The velocity can be selected from the change in arrival time as the source to geophone distance changes for simple situations.
In practice, the seismic reflection method is much more complicated. Reflections from most of the many interfaces within the Earth are very weak and so do not stand out against background noise. The reviews from closely spaced interfaces interfere with each other. Reflections from interfaces with different dips, seismic waves that repeatedly bounce between interfaces (“multiples”), converted waves, and waves traveling by other modes interfere with desired reflections. Also, velocity irregularities bend seismic rays in ways that are sometimes complicated.
The objective of most seismic work is to map geologic structures by determining the arrival time of reflectors. However, changes in the amplitude and wave shape contain information about stratigraphic changes and occasionally hydrocarbon accumulations. In some cases, seismic patterns can be identified with depositional systems, unconformities, channels, and other features.
The seismic reflection method usually gives better resolution (i.e., makes it possible to see more minor features) than other methods, except measurements made nearby, as with borehole logs. Thus, in most exploration programs, appreciably, more money is spent on seismic reflection work than on all other geophysical methods combined.
Fundamental aspects of Rock Mechanics
Engineering rock mechanics
The term engineering rock mechanics describes the engineering application of rock mechanics to civil, mining, petroleum, and environmental engineering circumstances. The term mechanics means studying the equilibrium and motion of bodies, which includes statics and dynamics. Thus, rock mechanics is the study of mechanics applied to rock and rock masses. ‘Engineering rock mechanics’ is this study within an engineering context, rather than in the context of natural processes that occur in the Earth’s crust, such as folding and faulting. Thus, the term rock engineering refers to the process of engineering with rock, and especially to creating structures on or in rock masses, such as slopes alongside roads and railways, dam foundations, shafts, tunnels, caverns, mines, and petroleum wellbores.
There is an essential distinction between ‘rock mechanics’ and ‘rock engineering.’ When ‘rock mechanics’ is studied in isolation, there is no specific engineering objective. For example, the potential collapse of a rock mass is neither good nor bad: it is just a mechanical fact. However, if the collapsing rock mass is in the roof of a civil engineering cavern, there is an adverse engineering connotation. Conversely, if the collapsing rock mass is part of a block caving system in mining (where the rock mass is intended to fail), there is a beneficial engineering connotation. In the civil engineering case, the integrity of the cavern is maintained if the rock mass in the roof does not collapse. In the mining engineering case, the integrity of the mining operation is maintained if the rock mass does collapse.
Hence, rock engineering applies a subjective element to rock mechanics because of the engineering objective. The significance of the rock mass behavior lies in the eye and brain of the engineer, not in the mechanics.
The distinction between ‘rock mechanics’ and ‘rock engineering’ illustrated in Fig. 1.1 is highlighted further in Fig. 1.2, which shows part of the concrete foundation illustrated in the Frontispiece. ‘Rock mechanics’ involves characterizing the intact rock strength and the geometry and mechanical properties of the natural fractures of the rock mass. These studies can be studied with other aspects of the rock mass properties such as rock stiffness and permeability without reference to a specific engineering function. However, when the studies take on a generic engineering direction, such as the structural analysis of foundations, we are in engineering rock mechanics. This is analogous to the term engineering geology, in which geology is studied, not in its entirety but aspects relevant to engineering.
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Figure 1.1. The distinction between ‘rock mechanics’ itself (a) and engineering applications of rock mechanics (b). In (a), F1…Fn is the boundary forces caused by rock weight and current tectonic activity. In (b), a tunnel is being constructed in a rock mass.
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Figure 1.2. The portion of the Frontispiece photograph illustrating loading of discontinuous rock mass by the concrete support of a multi-story car park, Jersey, UK.
‘Rock engineering’ is concerned with specific engineering circumstances: in this case (Fig. 1.2), the consequences of loading the rock mass via the concrete support. How many loads will the rock foundation support under these conditions? Will the support load cause the rock to slip on the pre-existing fractures? Is the stiffness of the concrete backing a significant parameter in these deliberations? If the rock mass is to be reinforced with rock bolts, where should these be installed? How many rock bolts should there be? At what orientation should they be installed? The photograph in Fig. 1.2 highlights all these issues.
Above the Frontispiece photograph, there are two acronyms:
Chile — Continuous, Homogeneous, Isotropic and Linearly Elastic;
Diane — Discontinuous, Inhomogeneous, Anisotropic, and Not-Elastic.
These refer to two ways of thinking about and modeling the rock mass. In the Chile case, we assume an ideal type of material that is not fractured, or if it is cracked, the fracturing can be incorporated into the elastic continuum properties. In the DIANE case, the nature of the natural rock mass is recognized, and we model accordingly, still often making gross approximations. Rock mechanics started with the Chile approach and has now developed techniques to enable the Diane approach. It is evident from Fig. 1.2 that a Diane approach is essential for this problem, using information about the orientation and strength of the rock fractures. However, both methods have their advantages and disadvantages, and the wise rock engineer will utilize each to maximal benefit according to the circumstances.
Modeling for rock mechanics and rock engineering should be based on ensuring that the relevant mechanisms and the governing parameters relating to the problem at hand have been identified. Then, the choice of modeling technique is based on the information required, e.g., ensuring an adequate foundation, as illustrated in Fig. 1.2.
Civil engineers have to deal with geotechnical matters where natural conditions remain unknown routinely, and inferences must be made based on observations and experience, with some assistance from laboratory testing. By contrast, the applied science of mechanics and structural engineering is based on deduction that gives definite results. These two aspects have to be considered when you try to understand what rock mechanics is and where an engineer has to assess the properties and strengths of the rock that he can use for foundations for structures.
Rock mechanics determines how a particular rock reacts when it is put to the use required by humanity for buildings, roads, bridges, dams, tunnels, and other civil engineering uses. It will assess the bearing capacity of the rock on the surface and how the force applied on the rock by the structures being built on it will affect the rock at various depths. Rock mechanics will determine the rock’s shear strength, which will allow the stone to resist the forces applied to it. Rock mechanics can also determine rock response when subjected to dynamic loading that may result from artificial applications or natural occurrences like earthquakes. The failure mechanism of rocks will allow engineers to counteract these so that the structures built on the rock are safe. Rock mechanics will also study the effect of defects in the. Rock mechanics will also allow engineers to decide how to protect slopes, the proper technique for tunneling, the strengths that can be expected from rock that functions as ballast for railway tracks or as the base for roads. The power of rock also plays a large part in aggregate used for concrete, making up most of the buildings nowadays.
Testing in rock mechanics
While laboratory testing for rocks does give extensive data for engineers to determine bearing capacities, shear strengths, permeability, and other concerns for designers, it is being acknowledged that rock mechanics benefits most from in situ testing of the rock and observation of geological conditions that can affect the way a rock behaves when subjected to loads and stresses. Engineers can then decide whether the stability of rocks and rock slopes could affect their building structures. Besides making the stone vulnerable to fracture, cavities present in the rock can also act as reservoirs for water and other solutions that can affect civil engineering structures. This becomes especially critical in the case of dams and tunnels where these cavities can affect the system’s stability, and in the case of tunnels, it can affect traffic going through them. Rock mechanics, along with geological studies, can decide whether rock slides or rock falls could occur and the measures that would be needed to prevent them.
The rock mechanics test is the basis for obtaining the mechanical parameters of rock and a necessary means for studying rock mechanics and engineering. In this paper, the uniaxial compression deformation test, Brazilian splitting test, and cornea pressure shear test are carried out for rocks in the Dajishan tungsten mine. As a result, the fundamental mechanical parameters such as uniaxial compressive strength, tensile strength, elastic modulus, Poisson’s ratio, and internal friction angle of the ore rock and surrounding rock are obtained. Meanwhile, damage characteristics of rock are deeply studied and analyzed under different experimental conditions. According to rock mechanics parameters obtained from indoor rock mechanics tests, three design schemes of stope structure parameters are optimized using the FLAC3D numerical simulation software. On the premise of ensuring the stability of the stope structure, the recovery rate of ore and the production capacity of the stope are taken into consideration. It is suggested that the second scheme should be adopted for mines (18 m for ore room and 7 m for ore pillar), which provides scientific guidance for the safe and efficient mining of mines.
Environmental Geology
When most people hear someone mention geology, the first image that pops into their minds is usually either someone wielding a rock hammer or drilling for oil. However, most people don’t realize that there is so much more to geology than that. Geology is the branch of science that deals with the Earth, its materials, and its processes. Environmental geology is the branch of geology concerned with the interactions between humans and the geologic environment. Environmental geology is essentially a way of applying geologic knowledge to identify, remediate, and hopefully prevent ecological problems from occurring due to people.
Environmental geologists must have a solid understanding of not only currently occurring geologic events but historical geologic events, such as past earthquakes and floods. This knowledge of the past is necessary because it helps them to get a better idea of what types of geologic events repeat themselves, with what frequency they might occur, and what types of damage occurred because of those events. This is different from what a paleontologist (someone who studies fossils) would do because environmental geologists are concerned with how the past relates to the present.
Likewise, environmental geologists can also attempt to protect people from environmental factors beyond their control (like suggesting that they not build a home in an active flood plain). Thus, ecological geology as a field is just as broadly reaching and engaging as geology is, with many potential phenomena and human-Earth interactions to research.
| Environmental geologists help to produce environmental hazard maps, like this lava hazard map. |
Importance of Environmental Geology
Environmental geology is a fundamentally important branch of science because it directly impacts every person on the planet every day. There is simply no way to avoid the environment around you.
The decisions that people, businesses, and governments make regarding the environment and environmental issues impact countless people, not merely the person or people who made the original decision. Therefore, the human impact of natural and artificial ecological problems is a significant ethical concern, making proper understanding of the science behind these issues more important.
When most people hear someone mention geology, the first image that pops into their minds is usually either someone wielding a rock hammer or drilling for oil. However, most people don’t realize that there is so much more to geology than that. Geology is the branch of science that deals with the Earth, its materials, and its processes. Environmental geology is the branch of geology concerned with the interactions between humans and the geologic environment. Environmental geology is essentially a way of applying geologic knowledge to identify, remediate, and hopefully prevent ecological problems from occurring due to people.
Environmental geologists must have a solid understanding of not only currently occurring geologic events but historical geologic events, such as past earthquakes and floods. This knowledge of the past is necessary because it helps them to get a better idea of what types of geologic events repeat themselves, with what frequency they might occur, and what types of damage occurred because of those events. This is different from what a paleontologist (someone who studies fossils) would do because environmental geologists are concerned with how the past relates to the present.
Likewise, environmental geologists can also attempt to protect people from environmental factors beyond their control (like suggesting that they not build a home in an active flood plain). Thus, ecological geology as a field is just as broadly reaching and engaging as geology is, with many potential phenomena and human-Earth interactions to research.
Importance of competence of sites by grouting
The advantages of grouting include:
, e. Used for slab jacking to lift or level distorted foundations
Structures of any kind need to be supported by the soil and bedrock beneath them. While this concept may seem obvious, the subsurface conditions are not. Therefore, ground improvement is often needed to increase the bearing capacity, provide settlement control for new construction, or remedy a current issue. Helical Drilling provides many ground improvement solutions, two of which are the grouting techniques compaction grouting and jet grouting.
Compaction grouting improves the strength and stiffness of the soil by high-pressure injection of a cementitious grout mix through a small-diameter casing at a pre-determined depth. Displacement and compaction occur as the container is gradually removed and the expanded grout column is built from the bottom up in lifts. While Helical’s compaction grouting can be applied to new construction projects, improving the ground beneath existing structures is also an excellent choice. The grout can be applied by injecting at inclined angles to reach beneath existing foundations drilling and even directly through existing floor slabs. Though this technology is typically used to improve the engineering properties of loose fills and native soils, it can be used to enhance many different soil types both above and below the water table.
Another grouting method Helical has in its arsenal is jet grout columns. Jet grouting is a method that involves injecting the grout material under very high pressures through high-velocity jets so that they hydraulically cut, erode, replace, and mix with the existing soil to form very uniform, high-strength, soil-cement columns. These can be installed at inclined angles to underpin existing structures or overlapped to create seepage barriers, cutoff walls, or excavation support. Like compaction grouting, jet grouting can be implemented both above and below the water table.
These grouting techniques can be tailored to various applications: from improved strength and stability to controlling seepage. Being a design-build firm, Helical can customize a grouted ground improvement solution to meet your specific needs.
Grout Curtains Grout curtains are constructed by injecting particulate or chemical grouts under pressure. The types of grout most commonly used are particulate grouts such as portland cement. Grout curtains reduce the permeability and increase the mechanical strength of the soils but can be three times more expensive than slurry walls. Because of the expense, grouting is best suited to seal unsound rock and situations where other barrier walls are impractical. In addition to cost considerations, some chemical grouts such as phenolic, acrylamide, and polyester are not often used or are not available because their toxicity requires special care in handling and for safeguards after implementation.
Microfine Cement Thick slurries can not penetrate fine cracks, and higher injection pressures would cause fracturing of ground foundations. However, because of the higher water requirements of microfine cement, the slurry remains fluid enough to flow into and penetrate fine sands and small cracks in the rock. These types of glue can treat finer-grained sands not possible to treat with portland cement alone. They are also used to stabilize waste plumes. They are composed of ground slag and portland cement mixed with large quantities of water or dispersants to become more fluid. Micro fines can develop early strength, and the thickening time is optimized with retarders.
Permeable Reactive Barrier PRB walls are passive treatment walls because their underground construction intercepts contaminated groundwater, and funnels flow through paths of reactive material or “gates.” As groundwater flows through the reactive material, chemical, biological or physical processes treat contaminants, transforming them into harmless byproducts. They can be constructed by excavation and backfill methods or, as in most cases, by biopolymer trenching. A narrow trench is excavated and filled with biodegradable slurry. Shoring or dewatering is unnecessary since the slurry acts as shoring by exerting hydraulic pressure against the trench walls. Sand, zero-valent metals, chelators, sorbents, or microbes are mixed at the proper ratios and usually trimmed into the excavation.
Answers:
1)B 2)A 3)C 4)D 5)B 6)A 7)B 8)C 9)C 10)C 11)A 12)B 13)A 14)D 15)C 16)A 17)A 18)B 19)C 20)C 21)C 22)D 23)B 24)B 25)D 26)A 27)C 28)A 29)B 30)C 31)A 32)B 33)B 34)D 35)C 36)A 37)b 38)C 39)A 40)A
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