9 Crustal Deformation and Earthquakes
At the end of this chapter, students should be able to:
- Differentiate between stress and strain
- Differentiate between brittle, ductile, and elastic deformation
- Identify the three major types of stress and their associated plate tectonics boundary
- Name different fold types
- Name different faults
- Understand the elastic rebound theory
- Describe how seismic waves are measured
- Understand earthquake magnitude and how it is quantified
- Identify areas of increased seismic hazard
- Determine the location of an epicenter
- Describe notable historical earthquakes
- Explain how humans can induce seismicity
Tectonic processes produce horizontal forces in the crust that cause pushing, pulling, and shearing stresses that deform the rock. Stresses created by tectonics, gravity, and igneous pluton emplacement cause deformation in rock. The type of deformation, which can be folds, fractures, and/or faults, depends on the setting, timing, and rock material.
9.1 Stress and Strain
Stress is the force exerted per unit area and strain is a material’s response to that force. Strain is deformation caused by stress. Strain in rocks can be represented as a change in rock volume and/or rock shape, as well as fracturing the rock. There are three types of stress: tensional, compressional, and shear . Tensional stress involves pulling something apart in opposite directions, stretching and thinning the material. Compressional stress involves things coming together and pushing on each other, thickening the material. Shear stress involves transverse movement of a material moving past each other, like a scissor.
|Type of Stress||Associated Plate Boundary type||Resulting Strain||Associated fault and offset types|
|Tensional||divergent||Stretching and thinning||Normal|
|Compressional||convergent||Shortening and thickening||Reverse|
When rocks are stressed, the resulting strain can be elastic, ductile, or brittle. This change is generally called deformation. Elastic deformation is strain that is reversible after a stress is released. For example, when you compress a spring, it elastically returns to its original shape after you release it. Ductile deformation occurs when enough stress is applied to a material that the changes in its shape are permanent, and the material is no longer able to revert to its original shape. For example, if you stretch a spring too far, it can be permanently bent out of shape. Note that concepts related to ductile deformation apply at the visible (macro) scale, and deformation is more complex at a microscopic scale. Research of plastic deformation, which touches on the atomic scale, is generally beyond the scope of introductory texts. Yield point is the amount of strain at which elastic deformation is surpassed and permanent deformation is measurable. In the figure, yield point is where the line transitions from elastic deformation to ductile deformation (the end of the dashed line). Brittle deformation is when the material undergoes another critical point of no return. When sufficient stress to pass that point occurs, it fails and fractures.
Important factors that influence if or how a rock will undergo elastic, ductile, or brittle deformation are intensity of the applied stress, time, temperature, confining pressure, pore pressure, strain rate, and rock strength. Pore pressure is the pressure exerted by fluids inside of the open spaces (pores) inside of a rock or sediment. Strain rate is how quickly a material is deformed. Rock strength is a measure of how easily a rock will respond to stress. Shale has low strength and granite has high strength.
Removing heat (decreasing temperature) makes the material more rigid. Likewise, heating materials make them more ductile. Heating glass makes it capable of bending and stretching. In terms of strain response, it is easier to bend a piece of wood slowly without breaking it.
|Increase Temperature||More Plastic|
|Increase Strain Rate||More Brittle|
|Increase Rock Strength||More Brittle|
9.3 Field Geology and Geological Maps
Topographic maps are two-dimensional (2D) representations of a three-dimensional (3D) land surface. Similarly, geologic maps are 2D representations of 3D geologic structures at the earth’s surface. Geologists use geologic maps to represent where geologic formations, faults, folds, and inclined rock units are. Geologic formations are recognizable, mappable rock units. In a geologic map, each formation drawn on the map is recognized by a color and an abbreviated label. For examples of geologic maps, check out the UGS geologic map viewer.
Symbols for formations on geologic maps have symbols formed in a specific way. The first capital letter(s) of that label represents the geologic time period of the formation, while the following lowercase letters represent the formation name and/or an abbreviated rock type description. Where more than one capital letter begins the symbol, it indicates multiple time periods for the formation.
9.3.1 Cross sections
Cross sections are subsurface interpretations from surface and subsurface measurements. Maps display geology in the horizontal plane, while cross sections show subsurface geology in the vertical plane. For more information on cross sections, check out the AAPG wiki on cross sections.
9.3.2 Strike and Dip
Geologists use a special symbol called strike and dip to represent beds that are inclined. Strike and dip symbols look like the capital letter “T” on a map with a wide top of the T. The short trunk of the “T” represents the dip direction of the inclined rock bed. Oftentimes, the dip symbol will have a number next to it that represents dip angle. Dip is the angle that a bed plunges into the Earth from the horizontal. One way to visualize strike is to think about a pitched roof on a rectangular house. The strike of the roof would be indicated by the horizontal line at the top of the roof or the eave that extends in a compass direction (NSEW). The strike is the angle between that horizontal line and true north or true south, e.g. N 43° E, meaning the horizontal line points toward the NE at an angle of 43° from true north. The dip of the roof would represent how steep the roof is with respect to horizontal. The direction of dip would be the same direction that a ball would roll off of the roof from stationary. A horizontal rock bed has a dip of 0°, and a vertical bed has a dip of 90°.
Geologic folds are layers of rock that are curved or bent by ductile deformation. Terms involved with folds include axis, which is the line along which the bending occurred, and limbs, which are the dipping beds that make up the sides of the folds. Folds are most commonly formed by compressional forces at depth, where hotter temperatures and higher confining pressures allow ductile deformation to occur.
Folds are described by the orientation of their axes, axial planes, and limbs. A fold is made up of two or more sets of dipping beds, generally dipping in opposite directions, that come together along a line, called the axis. Each set of dipping beds is known as a fold limb. The plane that splits the fold into two halves is known as the axial plane.
Symmetrical folds have mirrored limbs across their axial planes. The limbs of a symmetrical fold are inclined at the same (but opposite) angle indicating equal compression on both sides of the fold. Asymmetrical folds have dipping, non-vertical axial planes, where limbs dip into the ground at different angles. Recumbent folds are very tight folds with limbs compressed near the axial planes, and are generally horizontal, and overturned folds are where the angles on both limbs dip in the same direction. The fold axis is where the axial plane intersects the strata involved in the fold. A horizontal fold has a horizontal fold axis. When the axis of the fold plunges into the ground, the fold is called a plunging fold.
Anticlines are arch-like (“A”-shaped) folds, with downward curving limbs that have beds that dip away from the central axis of the fold. They are convex-upward in shape. In anticlines, the oldest rock strata are in the center of the fold, along the axis, and the younger beds are on the outside. An antiform has the same shape as an anticline, but in antiforms the relative ages of the beds in the fold cannot be determined. Oil geologists have interest in anticlines because they can form oil traps, where oil migrates up along the limbs of the fold and accumulates in the high point along the axis of the fold.
Synclines are trough-like (“U” shaped) , upward curving folds that have beds that dip in towards the central axis of the fold. They are concave-upward in shape. In synclines, the older rock is on the outside of the fold and the youngest rock is on the inside of the fold along the axis. A synform has the shape of a syncline but, like an antiform, does not distinguish between the ages of the units.
Monoclines are step-like folds, in which flat rocks are upwarped or downwarped, then continue flat. Monoclines are relatively common on the Colorado Plateau where they form “reefs,” which are ridges that act as topographic barriers and should not be confused with ocean reefs. Capitol Reef is an example of a monocline in Utah. Monoclines can be caused by bending of shallower sedimentary strata as faults grow below them. These faults are commonly called “blind faults” because they end before reaching the surface and can be either normal or reverse faults.
A dome is a symmetrical to semi-symmetrical upwarping of rock beds. Domes have a shape like an inverted bowl, similar to domes on buildings, like the Capitol Building. Domes in Utah include the San Rafael Swell, Harrisburg Junction Dome, and the Henry Mountains . Some domes are formed from compressional forces, while other domes are formed from underlying igneous intrusions , by salt diapirs, or even impacts, like upheaval dome in Canyonlands National Park.
A basin is the inverse of a dome. The basin is when rock forms a bowl-shaped depression. The Uinta Basin is an example of a basin in Utah. Technically, geologists refer to rocks folded into a bowl-shape as structural basins. Sometimes structural basins can also be sedimentary basins in which large quantities of sediment accumulate over time. Sedimentary basins can form as a result of folding, but are much more commonly produced in mountain building, between mountain blocks or via faulting. Regardless of the cause, as the basin sinks (called subsidence), it can accumulate even more sediment as the weight of the sediment causes more subsidence in a positive-feedback loop. There are active sedimentary basins all over the world . An example of a rapidly subsiding basin in Utah is the Oquirrh Basin of Pennsylvanian-Permian age in which over 30,000 feet of fossliferous sandstones, shales, and limestones accumulated. These strata can be seen in the Wasatch Mountains along the east side of Utah Valley, especially on Mt. Timpanogos and in Provo Canyon.
Faults are the places in the crust where brittle deformation occurs as two blocks of rocks move relative to one another. There are three major fault types: normal, reverse, and strike-slip. Normal and reverse faults display vertical, also known as dip-slip, motion. Dip-slip motion consists of relative up and down movement along a dipping fault between two blocks, the hanging wall and the footwall. In a dip-slip system, the footwall is below the fault plane and the hanging-wall is above the fault plane. A good way to remember this is to imagine a mine tunnel through a fault; the hanging wall would be where a miner would hang a lantern and the footwall would be at the miner’s feet. Faults are more prevalent near and related to plate boundaries, but can occur in plate interiors as well. Faults can show evidence of movement along the fault plane. Slickensides are polished, often grooved surfaces along the fault plane created by friction during the movement. A joint or fracture is a plane of breakage in a rock that does not show movement or offset. Joints can result from many processes, such as cooling, depressurizing, or folding. Joint systems may be regional affecting many square miles.
Normal faults move by a vertical motion where the hanging-wall moves downward relative to the footwall along the dip of the fault. Normal faults are created by tensional forces in the crust. Normal faults and tensional forces are commonly caused at divergent plate boundaries and where the crust is being stretched by tensional stresses. Utah examples of normal faults are the Wasatch Fault, the Hurricane Fault, and other faults bounding valleys in the Basin and Range.
Grabens, horsts, and half-grabens are all blocks of crust or rock that are bounded by normal faults. Grabens drop down relative to adjacent blocks and create valleys. Horsts go up relative to adjacent down-dropped blocks, and become areas of high topography. Where together, horsts and grabens create an symmetrical pattern of valleys surrounded by normal faults on both sides and mountains. Half-grabens are a one-sided version of a horst and graben, where blocks are tilted by a normal fault on one side, creating an asymmetrical valley-mountain arrangement. The mountain-valleys of the Basin and Range Province of Western Utah and Nevada consist of a series of full and half-grabens from the Salt Lake Valley to the Sierra Nevada Mountains. When the dip of a normal fault decreases with depth (i.e. the fault becomes more horizontal as it goes deeper), the fault is a listric fault. Extreme versions of listric faulting occur when large amounts of extension occur along very low-angle normal faults, known as detachment faults. The normal faults of the Basin and Range appear to become detachment faults at depth.
Reverse faults are when the hanging-wall moves up relative to the footwall. Reverse faults are caused by compressional forces. A thrust fault is a reverse fault where the fault plane has a low dip angle (generally less than 45 degrees). Thrust faults bring older rocks on top of younger rocks and can cause repetition of rock units in the stratigraphic record. Convergent plate boundaries with subduction zones create a special type of “reverse” fault called a megathrust fault. Megathrust faults cause the largest magnitude earthquakes and commonly cause tsunamis.
Strike-slip faults have side to side motion. In pure strike-slip motion, crustal blocks on either side of the fault do not move up or down relative to each other. There is left-lateral (sinistral) and right-lateral (dextral) strike slip motion. In left-lateral or sinistral strike slip motion, the opposite block moves left relative to the block that the observer is standing on. In right-lateral or dextral strike slip motion, the opposite block moves right relative to the observer’s block. Strike-slip faults are most commonly associated with transform boundaries, and are prevalent in fracture zones adjacent to mid-ocean ridges.
Bends in strike-slip faults can create areas where the sliding blocks create compression or tension. Tensional stresses will create transtensional features with normal faults and basins (like California’s Salton Sea), and compressional stresses will create transpressional features with reverse faults and small-scale mountain building (like California’s San Gabriel Mountains). The faults that splay off of transpression or transtension features are known as flower structures.
An example of a right-lateral strike-slip fault is the San Andreas Fault, which denotes a transform boundary between the North American and Pacific plates. An example of a left-lateral strike-slip fault is the Dead Sea fault in Jordan and Israel.
9.6 Earthquake Essentials
People feel approximately 1 million earthquakes a year. Few are noticed very far from the source. Even fewer are major earthquakes. Earthquakes are usually felt only when they are greater than a magnitude 2.5. The USGS Earthquakes Hazards Program has a realtime map showing the most recent earthquakes. Most earthquakes occur along active plate boundaries. Intraplate earthquakes (not along plate boundaries) are still poorly understood.
Earthquake energy is known as seismic energy, and it travels through the earth in the form of seismic waves. To understand some of the basics of earthquakes and how they are measured, consider some of the basic properties of waves. Waves describe a motion that repeats itself in a medium (rock or unconsolidated sediments in our case). The magnitude (height) of the motion is the amplitude of a wave. Wavelength is the distance between two successive peaks of the wave. The number of repetitions of the motion over time (cycles per time) is the frequency. The inverse of frequency, which is the amount of time for a wave to travel one wavelength, is the period. When multiple waves combine, they can interfere with each other. When the waves are in sync with each other, they will have constructive interference, where the influence of one wave will add to and magnify the other. If the waves are out of sync with each other, they will have destructive interference. If two waves have the same amplitude and frequency and they are ½ wavelength out of sync, the destructive interference between them can eliminate each wave. This process of constructive and destructive interference is illustrated below.
9.6.2 How Earthquakes Happen
The release of seismic energy is explained by the elastic rebound theory. When rock is strained to the point that it undergoes brittle deformation, built-up elastic energy is released during displacement, which in turn radiates away as seismic waves. When the brittle deformation occurs, it creates an offset between the fault blocks at a starting point called the focus. This offset propagates along the surface of rupture, which is known as the fault plane.
The fault blocks of persistent faults like the Wasatch Fault of Utah are locked together by friction. Over hundreds to thousands of years, stress builds up along the fault. Eventually, stress along the fault overcomes the frictional resistance, and slip initiates as the rocks break. The deformed rocks “snap back” toward their original position in a process called elastic rebound. Bending of the rocks near the fault may reflect this build-up of stress and in earthquake prone areas like California, strain gauges that measure this bending are set up in an attempt to understand more about predicting an earthquake. In some locations where the flt is not locked, seismic stress causes continuous movement along the fault called fault creep, where displacement occurs gradually. Fault creep occurs along some parts of the San Andreas Fault.
Release of seismic energy occurs in a series of steps. After a seismic energy release, energy begins to build again during a period of inactivity along the fault. The accumulated elastic strain may produce small earthquakes (on or near the main fault). These are called foreshocks and can occur hours or days before a large earthquake, but they may not occur at all. The main release of energy occurs during the major earthquake, known as the mainshock. Aftershocks may then occur to adjust strain that built up from movement of the fault. They generally decrease over time.
9.6.3 Focus and Epicenter
The focus (aka hypocenter) of an earthquake is the point of initial breaking or rupturing where displacement of rocks. The focus is always at some depth below the ground surface in the crust, and not at the surface. From the focus, the displacement propagates up, down, and laterally along the fault plane. The displacement produces shock waves, called seismic waves. Generally speaking, the larger the displacement and the further it propagates, the greater the amount of shaking produced. More shaking is usually the result of more seismic energy released. The epicenter is the location on the Earth’s surface vertically above the point of rupture (focus). This is the location that most news reports give because it is the center of the area where people are affected. The focus is the point along the fault plane from which the seismic waves spread outward.
9.6.4 Seismic Waves
Seismic waves are an expression of the energy released after an earthquake. Seismic waves occur as body waves and surface waves. When seismic energy is released, the first waves to propagate out are body waves that pass through the body of the planet. Body waves include primary waves (P waves) and secondary waves (S waves). Primary waves are the fastest seismic waves. They move through rock via compression, very much like sound waves move through air. Particles of rock move forward and back during passage of the P waves. Primary waves can travel through both fluids and solids. Secondary waves travel slower and follow primary waves, propagating as shear waves. Particles of rock move from side to side during passage of S waves. Because of this, secondary waves cannot travel through fluids, including liquids, plasma, or gas.
When an earthquake occurs at a location in the earth, the body waves radiate outward, passing through the earth and into the rock of the mantle as a sub-spherical wave front. A point on this spreading wave front travels along a specific path which reaches a seismograph located at one of thousands of seismic stations scattered over the earth. That specific travel path is a line called a seismic ray. Since the density (and seismic velocity) of the mantle increases with depth, a process called refraction causes earthquake rays to curve away from the vertical and bend back toward the surface, passing through bodies of rock along the way.
Surface waves are produced when P and S body waves strike the surface of the earth and travel along the Earth’s surface, radiating outward from the epicenter. Surface waves travel more slowly than body waves. They have complex horizontal and vertical ground movement that creates a rolling motion. Because they propagate at the surface and have complex motions, surface waves are responsible for most of the damage. Two types of surface waves are Love waves and Rayleigh waves. Love waves produce horizontal ground shaking and, ironically from their name, are the most destructive. Rayleigh waves produce an elliptical motion of points on the surface, with longitudinal dilation and compression, like ocean waves. However, with Raleigh waves rock particles move in a direction opposite to that of water particles in ocean waves.
Earth is like a bell, and an earthquake is a way to ring it. Like other waves, seismic waves bend and bounce when passing from one material to another, like moving from a dense rock to a rock with even higher density. When a wave bends as it moves into a different substance, it is known as refraction, and when waves bounce back, it is known as reflection. Because S waves cannot move through liquid, they are blocked by the liquid outer core, creating a shadow zone on the opposite side of the planet to the earthquake source.
9.7 Measuring Earthquakes
Seismographs are instruments used to measure seismic waves. They measure vibration of the ground using pendulums or springs. The principle of the seismograph involves mounting a recording device solidly to the earth and suspending a pen or writing instrument above it on a spring or pendulum. As the ground shakes, the suspended pen records the shaking on the recording device. The graph resulting from measurements of a seismograph is a seismogram. Seismographs of the early 20th century were essentially springs or pendulums with pens on them that wrote on a rotating drum of paper. Digital ones now use magnets and wire coils to measure ground motion. Typical seismograph arrays measure vibrations in three directions: north-south (x), east-west (y), and up-down (z).
To determine distance of the seismograph from the epicenter, seismologists use the difference between the times when the first P waves and S waves arrive. After an earthquake, P waves will appear first on the seismogram, followed by S waves, and finally body waves, which have the largest amplitude on the seismogram. Surface waves do lose energy quickly, so they are not measured at great distances from the focus. Seismographs across the globe record arrivals of waves from each earthquake at many station sites. The distance to the epicenter can be determined by comparing arrival times of the P and S waves. Electronic communication among seismic stations and connected computers used to make calculations mean that locations of earthquakes and news reports about them are generated quickly in the modern world.
9.7.2 Locating Earthquake Epicenters with Triangulation
Each seismograph gives the distance from that station to the earthquake epicenter. Three or more seismograph stations are needed to locate the epicenter of an earthquake through triangulation. Using the arrival-time difference from the first P wave to first S wave, one can determine the distance from the epicenter, but not the direction. The distance from the epicenter to each station can be plotted as a circle, the distance being equal to the circle’s radius. The place where the circles intersect demarks the epicenter. This method also works in three dimensions with spheres and multi-axis seismographs to locate not only the epicenter but also the depth of the focus of the earthquake.
9.7.3 Seismograph Network
The International Registry of Seismograph Stations lists more than 20,000 seismographs on the planet. Seismologists can use and compare data from sets of multiple seismometers dispersed over a wide area, which is a seismograph network. By collaborating, scientists can map the properties of the inside of the earth, detect detonation of large explosive devices, and predict tsunamis. The Global Seismograph Network, a set of world-wide linked seismographs that distribute real time data electronically, consists of more than 150 stations that meet specific design and precision standards. The Global Seismograph Network helps the Comprehensive Nuclear-Test-Ban Treaty Organization monitor for nuclear tests. The USArray is a network hundreds of permanent and transportable seismographs in the United States. The USArray is being used to map the subsurface through passive collection of seismic waves created by earthquakes (see below).
Nepal Earthquake M7.9 Ground Motion Visualization
9.7.4 Seismic tomography
Very much like a CT (Computed Tomography) scan uses X-rays at different angles to image the inside of a body, seismic tomography uses rays from seismic waves created by thousands of earthquakes that occur each year and pass at all angles through masses of rock within the earth to generate images of internal structures.
Based on the assumption that the earth consists of homogenous layers, geologists have developed a model of expected properties of earth materials at every depth within the earth called the PREM (Preliminary Reference Earth Model). Included in these expected properties is transmission velocity of seismic waves which is dependent on density and elasticity of the rock. In the mantle, density differences in rock bodies result primarily from differences in temperature. Slightly cooler rocks have a higher density and therefore transmit earthquake waves slightly faster than the velocity predicted by PREM. Slightly warmer rocks transmit earthquake waves slightly slower than predicted by PREM. These small differences from PREM are called seismic anomalies and can be measured for bodies of rock within the earth from arrival times of seismic rays passing through them at stations of the seismic network. Such seismically defined bodies of rock can thus be imaged via seismic tomography by the network of seismic stations distributed over the earth.
Seismograph networks provide data for creating tomographic images and maps of the distribution of rock density beneath the crust. For example, seismologists have mapped the Farallon Plate, a tectonic plate that subducted beneath North America during the last several million years, and the Yellowstone magma chamber, which is a product of the Yellowstone hot spot under the North American continent. Peculiarities of the subduction of the Farallon Plate are thought to be responsible for many features of western North America including the Rocky Mountains (See chapter 8 for more information on the Farallon Plate).
9.7.5 Determining Earthquake Magnitude
Magnitude is the measure of intensity of an earthquake. The Richter scale is the most well-known magnitude scale devised for an earthquake, and was the first one developed by Charles Richter at CalTech. This was the magnitude scale used historically by early seismologists. The Richter scale magnitude is determined from measurements on a seismogram. Magnitudes on the Richter scale are based on measurements of the maximum amplitude of the needle trace measured on the seismogram and the arrival time difference of S and P waves which gives the distance to the earthquake.
The Richter scale is a logarithmic scale, based on powers of 10. Amplitude of the seismic wave recorded on the seismogram is 10 times greater for each increase of 1 unit on the Richter scale. That means a magnitude 6 earthquake shakes the ground 10 times more than a magnitude 5. However the actual energy released for each 1 unit magnitude increase is 32 times greater. That means energy released for a magnitude 6 earthquake is 32 times greater than a magnitude 5. The Richter scale was developed for distances appropriate for earthquakes in Southern California and on seismograph machines in use there. Its applications to larger distances and very large earthquakes is limited. Therefore, most agencies no longer use the methods of Richter to determine magnitude, but generate a quantity called the Moment Magnitude, which is more accurate for large earthquakes measured at the seismic array across the earth. As numbers, the moment magnitudes are comparable to the magnitudes of the Richter Scale. The media still often give magnitudes as Richter Magnitude even though the actual calculation was of moment magnitude.
9.7.6 Moment magnitude scale
The Moment Magnitude scale depicts the absolute size of earthquakes, comparing information from multiple locations and using a measurement of actual energy released calculated from cross-sectional area of rupture, amount of slippage, and the rigidity of the rocks. Because of the unique geologic setting of each earthquake and because rupture area is often hard to measure, estimates of moment magnitude can take days to months to calculate.
Like Richter magnitude, the moment magnitude scale is logarithmic. Both scales are used in tandem because the estimates of magnitude may change after a quake. The Richter scale is used as a quick determination immediately following the quake (and thus is usually reported in news accounts), and the moment magnitude is calculated days to months later. Magnitude values of the two magnitudes are approximately equal except for very large earthquakes.
9.7.7 Modified Mercalli Intensity Scale
The Modified Mercalli Intensity Scale is a qualitative scale (I-XII) of the intensity of ground shaking based on damage to structures and people’s perceptions. This scale can vary depending on the location and population density (urban vs rural). It was also used for historic earthquakes which occurred before quantitative measurements of magnitude could be made. The Modified Mercalli Intensity maps show where the damage is most severe based on questionnaires sent to residents, newspaper articles, and reports from assessment teams. Recently, USGS has used the internet to help gather data more quickly.
|I||Not felt||Not felt except by a very few under especially favorable conditions.|
|II||Weak||Felt only by a few persons at rest,especially on upper floors of buildings.|
|III||Weak||Felt quite noticeably by persons indoors, especially on upper floors of buildings.
Many people do not recognize it as an earthquake. Standing motor cars may rock slightly. Vibrations similar to the passing of a truck. Duration estimated.
|IV||Light||Felt indoors by many, outdoors by few during the day. At night, some awakened.
Dishes, windows, doors disturbed; walls make cracking sound. Sensation like heavy truck striking building. Standing motor cars rocked noticeably.
|V||Moderate||Felt by nearly everyone; many awakened. Some dishes, windows broken. Unstable objects overturned. Pendulum clocks may stop.|
|VI||Strong||Felt by all, many frightened. Some heavy furniture moved; a few instances of fallen plaster. Damage slight.|
|VII||Very strong||Damage negligible in buildings of good design and construction; slight to moderate in well-built ordinary structures; considerable damage in poorly built or badly designed structures; some chimneys broken.|
|VIII||Severe||Damage slight in specially designed structures; considerable damage in ordinary substantial buildings with partial collapse. Damage great in poorly built structures. Fall of chimneys, factory stacks, columns, monuments, walls. Heavy furniture overturned.|
|IX||Violent||Damage considerable in specially designed structures; well-designed frame structures thrown out of plumb. Damage great in substantial buildings, with partial collapse. Buildings shifted off foundations.|
|X||Extreme||Some well-built wooden structures destroyed; most masonry and frame structures destroyed with foundations. Rails bent.|
Shake maps (written ShakeMaps by the USGS) use high-quality seismograph data from seismic networks to show areas of intense shaking. They are the result of rapid, computer-interpolated seismograph data. They are useful in crucial minutes after an earthquake, as they can show emergency personnel where the greatest damage likely occurred and locate areas of possible damaged gas lines and other utilities.
9.8 Earthquake Risk
9.8.1 What determines shaking?
In general, the larger the magnitude, the stronger the shaking and the longer the shaking will last. But, other factors influence the level of shaking as described in the following paragraphs.
Table and descriptions from https://earthquake.usgs.gov/learn/topics/mag_vs_int.php
|Magnitude||Modified Mercalli Intensity||Shaking/Damage Description|
|1.0 – 3.0||I||Only felt by a very few.|
|3.0 – 3.9||II – III||Noticeable indoors, especially on upper floors.|
|4.0 – 4.9||IV – V||Most to all feel it. Dishes, doors, cars shake and possibly break.|
|5.0 – 5.9||VI – VII||Everyone feels it. Some items knocked over or broken. Building damage possible.|
|6.0 – 6.9||VII – IX||Frightening amounts of shaking. Significant damage especially with poorly constructed buildings|
|≥ 7.0||≥ VIII||Significant destruction of buildings. Potential for objects to be thrown in air from shaking.|
Location and Direction
Closer earthquakes will inherently cause more shaking than those farther away. The location in relation to epicenter and direction of rupture will influence how much shaking is felt. The direction that the rupture propagates along the fault influences the shaking. The path of greatest rupture can intensify shaking in an effect known as directivity.
Local Geologic conditions
The nature of the ground materials affects the properties of the seismic waves. Different materials respond differently to an earthquake. Think of shaking jello versus shaking a meatloaf, one will jiggle much more to the same amount of shaking. The response to shaking depends on their degree of consolidation; lithified sedimentary rocks and crystalline rocks shake less than unconsolidated sediments and land fill.
This is because seismic waves move faster through consolidated bedrock, move slower through unconsolidated sediment, and move slowest through unconsolidated materials with high water content. Since the energy is carried by both velocity and amplitude, when a seismic wave slows down, its amplitude increases, which in turn increases seismic shaking. Energy is transferred to the vertical motion of the surface waves.
Depth of focus
The focus is the place within the Earth where the earthquake starts. The depth of earthquakes influences the amount of shaking. Deeper earthquakes cause less shaking at the surface because they lose much of their energy before reaching the surface. Recall that most of the destruction is caused by surface waves which are caused as the body waves reach the surface.
9.8.2 What determines destruction?
Building material choices can influence the amount of damage caused by earthquake shaking. The flexibility of building materials relates to their resistance to damage by earthquake waves. Unreinforced Masonry (URM) is the most devastated by ground shaking. Wood framing held together with nails which can bend and flex with wave passage are more likely to survive earthquakes. Steel also has the ability to deform elastically before brittle failure. The Salt Lake City campaign “Fix the Bricks” has good information on URMs and earthquake safety.
Shaking Intensity and Duration
Greater shaking and duration of shaking will cause more destruction than less shaking and shorter shaking.
Resonance is when the frequency of seismic energy matches a building’s natural frequency of shaking, determined by properties of the building, and intensifies the amplitude of shaking. This famously happened in the 1985 Mexico City Earthquake, where buildings of heights between 6 and 15 stories were especially vulnerable to earthquake damage. Skyscrapers designed with earthquake resilience have dampers and base isolation features to reduce resonance.
9.8.3 Earthquake Recurrence
Geologists dig earthquake trenches across some faults to measure ground deformation and estimate the frequency of occurrence of past earthquakes. Trenches are effective for faults with relatively long recurrence intervals (100s to 10,000s of years), which is the period of time between significant earthquakes. In areas with more frequent earthquakes and more measured earthquake data, trenches are less necessary. A long hiatus in earthquake activity could indicate the buildup of stress on a specific segment of a fault with strain held in place by friction, which would indicate a higher probability of an earthquake along that segment. This hiatus of seismic activity along a length of a fault (i.e. a fault that is locked and not having any earthquakes) is known as seismic gap.
9.8.4 Distribution of Earthquake Hazard
Subduction zones are where the largest, deepest earthquakes occur. These are known as megathrust earthquakes. Example areas include the Sumatran Islands, the Aleutian Islands, and the west coast of South America. The Cascadia Subduction Zone off the coast of Washington and Oregon is another exmple.
Collision zone earthquakes
Continental collisions create broad area of earthquakes. They can have some deep, large earthquakes from ‘left-over’ subduction and/or deep-crustal processes. An area where this is occurring is the Himalayan Mountains and the Alps.
Transform Fault Boundaries
Transform fault boundaries create moderate and large earthquakes, usually having a maximum magnitude of about 8. The San Andreas fault in California is an example of a transform fault boundary. Other examples are the Alpine Fault in New Zealand and the Anatolian Faults in Turkey.
Rifts and mid-ocean ridges
Continental rifts and mid-ocean ridges are characteristic of divergent plate boundaries. These areas generally produce moderate earthquakes. Examples of such areas include the East African Rift and Iceland. The United State’s Basin and Range is another area undergoing tensional forces that experiences earthquakes.
Intraplate earthquakes are earthquakes not near tectonic plate boundaries. Intraplate earthquakes generally occur in areas of weakened crust or concentrated tectonic stress. The New Madrid Seismic zone is an area in Missouri, Illinois, Tennessee, Arkansas, and Indiana that is thought to represent the failed Reelfoot Rift zone . The failed rift zone created an area of crustal weakness, which is relatively responsive to tectonic stresses related to plate movement and interaction. The infrequent earthquakes could be related to reactivated areas of weakness with a low rate of strain.
9.8.5 Secondary Hazards Caused by Earthquakes
Liquefaction is when saturated unconsolidated sediments (usually silt or sand) is liquefied from shaking. Shaking causes loss of cohesion between grains of sediment, reducing the effective stress resistance of the sediment. The sediment flows very much like the quicksand presented in movies. Liquefaction creates sand volcanoes, which is when liquefied sand is squirted through an overlying (usually finer-grained) layer, creating cone-shaped sand features. It may also cause buildings to settle or tilt.
Many of the more recent devastating natural disasters have been caused by earthquake-induced tsunamis. Tsunamis form when the sea floor is offset by earthquakes in the ocean subsurface. This offset can be caused by fault movement or underwater landslides and actually lifts a volume of ocean water generating the tsunami wave. Tsunami waves travel fast with low amplitude in deep ocean water, but are significantly amplified as the water shallows as they approach shore. When a tsunami is about to strike land, the water in front of the wave along the shore will recede significantly, tragically causing curious people to wander out. This receding water is the drawback of the trough in front of the tsunami wave which then crashes on shore as a wall of water upwards of a hundred feet high. The behavior of tsunamis as ocean waves is covered in the section on shorelines in Chapter 12. Warning systems have been established to help mitigate the loss of life caused by tsunamis.
Shaking can trigger landslides (see landslide section for more information). One example is the 1992 magnitude 5.9 earthquake in St. George Utah . This earthquake caused the Springdale landslide, having a scarp that offset and destroyed several structures in the Balanced Rock Hills subdivision .
Seiches are waves on lakes generated by earthquakes, which cause sloshing of water back and forth and, sometimes, even changes in elevation of the lake. A seich in Hebgen Lake during the 1959 earthquake caused major destruction to structures and roads around the lake.
Land Elevation Changes
Significant subsidence and upheaval of the land can occur in relation to the slippage that causes earthquakes. Land elevation changes are the result of the relaxation of stress and subsequent movement along the fault plane. The 1964 Alaska earthquake is a good example of this. Where the fault cuts the surface, elevation of one side causes a fault scarp that may be a few feet to 20 or 30 feet in height. The Wasatch Mountains represent an accumulation of fault scarps of a couple dozen feet at a time over a few million years.
9.8.6 Human Activity that can Create Seismic Energy
We have used seismographs to determine the size of nuclear weapons tested by other countries, most recently in North Korea . The Earthquake magnitude energy calculator can inform readers on the amount of energy that different magnitude earthquakes can produce.
9.9 Case Studies
9.9.1 Basin and Range Earthquakes
Basin and Range earthquakes are caused primarily by normal faults created by tensional forces pulling the area apart. The Wasatch Fault defines the eastern extent of the Basin and Range and has been studied as an earthquake hazard for more than 100 years. The Basin and Range extends from the Wasatch Fault to the Sierra Nevada.
9.9.2 North American Earthquakes
- 1811-1812 New Madrid Earthquakes – Historical accounts of the New Madrid seismic zone date as far back as 1699 . A sequence of large (moment magnitude >7) occurred from December 1811 to February 1812 in the New Madrid Missouri area . The earthquakes damaged houses in St. Louis, affected the stream course of the Mississippi River, and leveled the town of New Madrid. These earthquakes were the result of seismic activity in the New Madrid seismic zone, an area of intraplate seismic activity. The intraplate activity is thought to be derived from a failed Reelfoot rift zone (an aulacogen), creating crustal weakness in the region . The New Madrid seismic zone continues to produce earthquakes.
- 1868 Charleston – The 1868 earthquake of Charleston South Carolina was a moment magnitude 7.0, with a Mercalli intensity of X, killing at least 60 people. This was an intraplate earthquake, likely associated with ancient faults created during the breakup of Pangea. The earthquake caused significant liquefaction . Scientists estimate that destructive earthquakes may recur in this area with an interval of approximately 1500 to 1800 years.
- 1906 Great San Francisco Earthquake and Fire – On April 17, 1906, a large earthquake occurred along the San Andreas fault near San Francisco. The earthquake had an estimated moment magnitude of 7.8 and a Modified Mercalli Intensity of XI. Geologist G.K. Gilbert was present to take measurements and photographs after the earthquake . There were multiple aftershocks followed by fires that devastated the city. About 80% of the city was destroyed.
- 1964 Alaska – Magnitude 9.2 earthquake created by the megathrust fault along the Aleutian subduction zone. Large areas of land dropped down while other areas uplifted. The earthquake caused significant mass wasting (see landslides section). The 1964 Alaska earthquake was one of the most powerful earthquakes ever recorded.
- 1989 Loma Prieta – The Loma Prieta earthquake was a moment magnitude 6.9 earthquake created by movement along the San Andreas Fault. It caused 63 deaths and buckled portions of the freeway and part of the Oakland San Francisco Bay Bridge.
9.9.3 Global Earthquakes
- 1556 Shaanxi – On January 23, 1556 an earthquake of approximate magnitude 8 hit central China, killing approximately 830,000 people. This system is thought to have a recurrence interval of 1000 years . Much of the death toll was attributed to the collapse of loess (windblown sediment in an area in China) cave dwellings (yaodongs) that collapsed due to the shaking. This earthquake is considered the most deadly earthquake in history.
- 1755 Lisbon – On November 1, 1755 an earthquake with an estimated magnitude of 8-9 struck Lisbon Portugal , killing between 10,000 to 17,400 people . The earthquake was followed by a tsunami.
- 1960 Valdivia Chile – The 1960 Valdivia earthquake was the most powerful earthquake ever measured, having a moment magnitude between 9.4-9.6. The earthquake, occurring on May 22, is estimated to have lasted 10 minutes. It triggered a tsunami that destroyed houses in Japan and Hawaii, and caused a volcanic eruption of vents along the Cordón Caulle volcano.
- 1976 Tangshan earthquake – Just before 4 am on July 28, 1976 a magnitude 7.8 struck Tangshan, Hebei, China. This earthquake killed more than 240,000 people. The high death-toll is thought to be contributed to by the earthquake occurring early in the morning and building techniques that were not appropriate for earthquakes.
- 2004 Indonesia – On December 26, 2004, a moment magnitude 9.0-9.3 earthquake occurred off the coast of Sumatra, Indonesia . The earthquake was created by slippage of the Sunda Megathrust, where the Australia plate is subducted below the Sunda plate in the Indian Ocean (“Long-Term Perspectives on Giant Earthquakes and Tsunamis at Subduction Zones,” 2007). The earthquake resulted in a massive tsunami that is estimated to have killed over 200,000 people along the coastlines of the Indian Ocean, creating waves as tall as 24 meters when they reached the shore.
- 2010 Haiti – The magnitude 7 2010 Haiti earthquake occurred on January 12, 2010. It had many significant aftershocks at magnitude 4.5 or higher. It killed more than 92,000 people. Death toll was increased by destroyed and damaged infrastructure, which contributed to a cholera outbreak, among other issues.
- 2011 Tōhoku Japan – On March 11, 2011, Japan experienced a magnitude 9.0 earthquake. Because most of the buildings in Japan were designed to tolerate earthquakes, the earthquake caused a lot less damage than the tsunami it created. The tsunami caused tens of billions of dollars in damage and caused more than 15,000 deaths. The tsunami resulted in the meltdown and destruction of the Fukushima nuclear power plant.
9.9.4 Induced Seismicity
Injection of waste fluids in the ground, commonly a byproduct of an extraction process for natural gas known as fracking, can increase the outward pressure that liquid in the pores of a rock exerts, known as pore pressure . The increase in pore pressure decreases the frictional forces that keep rocks from sliding past each other, essentially lubricating fault planes. This effect is causing earthquakes to occur near injection sites, in a human induced activity known as induced seismicity . The significant increase in drilling activity in the central United States has spurred the requirement for the disposal of significant amounts of waste drilling fluid, resulting in a measurable change in the cumulative number of earthquakes experienced in the region.
Stress can come in the form of tension, shear, and compression which can generate normal, strike-slip, and reverse faults. Seismic energy is released when faults slip, and that energy can be measured and used to map the locations of earthquakes, the distribution of shaking, and the internal structures of our planet. When rock deformation is ductile instead of brittle, rocks can fold instead of faulting.