earthquake

Earthquakes Earthquakes occur when energy stored in elastically strained rocks is suddenly released. This release of energy causes intense ground shaking in the area near the source of the earthquake and sends waves of elastic energy, called seismic waves, throughout the Earth. Earthquakes can be generated by bomb blasts, volcanic eruptions, sudden volume changes in minerals, and sudden slippage along faults. Earthquakes are definitely a geologic hazard for those living in earthquake prone areas, but the seismic waves generated by earthquakes are invaluable for studying the interior of the Earth.It can also be defined as the rumbling or trembling of the ground produced by the sudden breaking of rocks in response to geological forces within the Earth In or discussion of earthquake we want to answer the following questions: What causes earthquakes? How are earthquakes studied? What happens during an earthquake? Where do earthquakes occur? Can earthquakes be predicted? Can humans be protected from earthquakes? What can earthquakes tell us about the interior of the earth? Causes of Earthquakes
Within the Earth rocks are constantly subjected to forces that tend to bend, twist, or fracture them. When rocks bend, twist or fracture they are said to deform.  Strain is a change in shape, size, or volume. The forces that cause deformation are referred to as stresses.  To understand the causes of earthquakes we must first explore stress and strain. Stress and Strain Recall that stress is a force applied over an area. A uniform stress is where the forces act equally from all directions. Pressure is a uniform stress and is referred and is also called confining stress or hydrostatic stress. If stress is not equal from all directions then the stress is a differential stress.
Three kinds of differential stress occur. Tensional stress (or extensional stress), which stretches rock; Compressional stress, which squeezes rock; and Shear stress, which result in slippage and translation.
When a rock is subjected to increasing stress it changes its shape, size or volume. Such a change in shape, size or volume is referred to as strain.  When stress is applied to rock, the rock passes through 3 successive stages of deformation.

Elastic Deformation -- wherein the strain is reversible. Ductile Deformation -- wherein the strain is irreversible. Fracture -- irreversible strain wherein the material breaks.
We can divide materials into two classes that depend on their relative behavior under stress. Brittle materials have a small to large region of elastic behavior, but only a small region of ductile behavior before they fracture. Ductile materials have a small region of elastic behavior and a large region of ductile behavior before they fracture.
How a material behaves will depend on several factors. Among them are: Temperature - At high temperature molecules and their bonds can stretch and move, thus materials will behave in more ductile manner. At low Temperature, materials are brittle. Confining Pressure - At high confining pressure materials are less likely to fracture because the pressure of the surroundings tends to hinder the formation of fractures. At low confining stress, material will be brittle and tend to fracture sooner. Strain rate -- Strain rate refers to the rate at which the deformation occurs (strain divided by time). At high strain rates material tends to fracture. At low strain rates more time is available for individual atoms to move and therefore ductile behavior is favored. Composition -- Some minerals, like quartz, olivine, and feldspars are very brittle. Others, like clay minerals, micas, and calcite are more ductile This is due to the chemical bond types that hold them together. Thus, the mineralogical composition of the rock will be a factor in determining the deformational behavior of the rock. Another aspect is presence or absence of water. In general, rocks near the surface of the earth behave in a brittle fashion, unless they are deformed slowly.   Thus, when they are acted upon by differential stress, they tend to fracture.

Faults Most natural earthquakes are caused by sudden slippage along a fault.  Faults occur when brittle rocks fracture and there is displacement of one side of the fracture relative to the other side.  The amount of displacement in a single slippage event is rarely more that 10 to 20 m for large earthquakes, but after many events the displacement could be several hundred kilometers. Types of Faults Faults can be divided into several different types depending on the direction of relative displacement or slip on the fault. Most faults make an angle with the ground surface, and this angle is called the dip angle.  If the dip angle is 90o the fault plane is vertical.  Faults can be divided into two major classes.  Dip Slip Faults - Dip slip faults are faults that have an inclined fault plane and along which the relative displacement or offset has occurred along the dip direction. Note that in looking at the displacement on any fault we don't know which side actually moved or if both sides moved, all we can determine is the relative sense of motion. For any inclined fault plane we define the block above the fault as the hanging wall block and the block below the fault as the footwall block
Normal Faults - are faults that result from horizontal extensional stresses in brittle rocks and where the hanging-wall block has moved down relative to the footwall block.
Reverse Faults- are faults that result from horizontal compressional stresses in brittle rocks, where the hanging-wall block has moved uprelative the footwall block.
A Thrust Faultis a special case of a reverse fault where the dip of the fault is less than 45o. Thrust faults can have considerable displacement, measuring hundreds of kilometers, and can result in older strata overlying younger strata.
Strike Slip Faults - are faults where the displacement on the fault has taken place along a horizontal direction. Such faults result from shear stresses acting in the crust. Strike slip faults can be of two varieties, depending on the sense of displacement. To an observer standing on one side of the fault and looking across the fault, if the block on the other side has moved to the left, we say that the fault is a left-lateral strike-slip fault. If the block on the other side has moved to the right, we say that the fault is a right-lateral strike-slip fault. The famous San Andreas Fault in California is an example of a right-lateral strike-slip fault. Displacements on the San Andreas fault are estimated at over 600 km.
Oblique Slip Faults - If the displacement has both a vertical component and a horizontal component (i.e. a combination of dip slip and strike slip) it is called an oblique slip fault. Blind Faults  A blind fault is one that does not break the surface of the earth.  Instead, rocks above the fault have behaved in ductile fashion and folded over the tip of the fault.

Active Faults An active fault is one that has shown recent displacement and likely has the potential to produce earthquakes.   Since faulting is part of the deformation process, ancient faults can be found anywhere that deformation has taken place in the past.  Thus, not every fault one sees is necessarily an active fault.  Surface Expression of Faults Where faults have broken the surface of the earth they can be delineated on maps and are called fault lines or fault zones.   Recent ruptures of dip slip faults at the surface show a cliff that is called a fault scarp.  Strike slip faults result in features like linear valleys, offset surface features (roads, stream channels, fences, etc.) or elongated ridges.(see figure 10.5 and10.33 in your textbook).

How Faults Develop
The elastic rebound theory suggests that if slippage along a fault is hindered such that elastic strain energy builds up in the deforming rocks on either side of the fault, when the slippage does occur, the energy released causes an earthquake.
This theory was discovered by making measurements at a number of points across a fault. Prior to an earthquake it was noted that the rocks adjacent to the fault were bending. These bends disappeared after an earthquake suggesting that the energy stored in bending the rocks was suddenly released during the earthquake.
Friction between the blocks then keeps the fault from moving again until enough strain has accumulated along the fault zone to overcome the friction and generate another earthquake.  Once a fault forms, it becomes a zone of weakness in the crust, and so long as the tectonic stresses continue to be present more earthquakes are likely to occur on the fault. Thus faults move in spurts and this behavior is referred to asStick Slip.  If the displacement during an earthquake is large, a large earthquake will be generated.  Smaller displacements generate smaller earthquakes.  Note that even for small displacements of only a millimeter per year, after 1 million years, the fault will accumulate 1 km of displacement. Fault Creep - Some faults or parts of faults move continuously without generating earthquakes.  This could occur if there is little friction on the fault and tectonic stresses are large enough to move the blocks in opposite directions.  This is called fault creep.  Note that if creep is occurring on one part of a fault, it is likely causing strain to build on other parts of the fault.

How Earthquakes Are Measured When an earthquake occurs, the elastic energy is released and sends out vibrations that travel in all directions throughout the Earth. These vibrations are called seismic waves.
The point within the earth where the fault rupture starts is called thefocus or hypocenter.  This is the exact location within the earth were seismic waves are generated by sudden release of stored elastic energy. The epicenter is the point on the surface of the earth directly above the focus. Sometimes the media get these two terms confused.

Seismic Waves Seismic waves emanating from the focus can travel in several ways, and thus there are several different kinds of seismic waves.

Body Waves - emanate from the focus and travel in all directions through the body of the Earth. There are two types of body waves: P-waves and S waves.

P - waves - are Primary waves. They travel with a velocity that depends on the elastic properties of the rock through which they travel.                                      Where, Vp is the velocity of the P-wave, K is the incompressibility of the material, μis the rigidity of the material, and ρ is the density of the material. P-waves are the same thing as sound waves. They move through the material by compressing it, but after it has been compressed it expands, so that the wave moves by compressing and expanding the material as it travels. Thus the velocity of the P-wave depends on how easily the material can be compressed (the incompressibility), how rigid the material is (the rigidity), and the density of the material. P-waves have the highest velocity of all seismic waves and thus will reach all seismographs first. S-Waves - Secondary waves, also called shear waves. They travel with a velocity that depends only on the rigidity and density of the material through which they travel:                                             S-waves travel through material by shearing it or changing its shape in the direction perpendicular to the direction of travel. The resistance to shearing of a material is the property called the rigidity. It is notable that liquids have no rigidity, so that the velocity of an S-wave is zero in a liquid. (This point will become important later). Note that S-waves travel slower than P-waves, so they will reach a seismograph after the P-wave.

Surface Waves - Surface waves differ from body waves in that they do not travel through the earth, but instead travel along paths nearly parallel to the surface of the earth. Surface waves behave like S-waves in that they cause up and down and side to side movement as they pass, but they travel slower than S-waves and do not travel through the body of the Earth.  Love waves result in side to side motion and Rayleigh waves result in an up and down rolling motion.  (see figure 10.10 in your text).  Surface waves are responsible for much of the shaking that occurs during an earthquake.

The study of how seismic waves behave in the Earth is called seismology. Seismic waves are measured and recorded on instruments called seismometers.
Seismometers
Seismic waves travel through the earth as elastic vibrations. Aseismometeris an instrument used to record these vibrations and the resulting graph that shows the vibrations is called aseismogram.
The seismometer must be able to move with the vibrations, yet part of it must remain nearly stationary. This is accomplished by isolating the recording device (like a pen) from the rest of the Earth using the principal of inertia. For example, if the pen is attached to a large mass suspended by a spring, the spring and the large mass move less than the paper which is attached to the Earth, and on which the record of the vibrations is made.

The record of an earthquake, a seismogram, as recorded by a seismometer, will be a plot of vibrations versus time. On the seismogram time is marked at regular intervals, so that we can determine the time of arrival of the first P-wave and the time of arrival of the first S-wave.

(Note again, that because P-waves have a higher velocity than S-waves, the P-waves arrive at the seismographic station before the S-waves).

Locating the Epicenter of an Earthquake

In order to determine the location of an earthquake, we need to have recorded a seismogram of the earthquake from at least three seismographic stations at different distances from the epicenter. In addition, we need one further piece of information - that is the time it takes for P-waves and S-waves to travel through the earth and arrive at a seismographic station. Such information has been collected over the last 100 or so years, and is available as travel time curves.

From the seismographs at each station one determines the S-P interval (the difference in the time of arrival of the first S-wave and the time of arrival of the first P-wave. Note that on the travel time curves, the S-P interval increases with increasing distance from the epicenter. Thus the S-P interval tells us the distance to the epicenter from the seismographic station where the earthquake was recorded.

Thus, at each station we can draw a circle on a map that has a radius equal to the distance from the epicenter. Three such circles will intersect in a point that locates the epicenter of the earthquake.

Earthquake Size  Whenever a large destructive earthquake occurs in the world the press immediately wants to know where the earthquake occurred and how big the earthquake was (in California the question is usually - Was this the Big One?). The size of an earthquake is usually given in terms of a scale called the Richter Magnitude. Richter Magnitude is a scale of earthquake size developed by a seismologist named Charles F. Richter. The Richter Magnitude involves measuring the amplitude (height) of the largest recorded wave at a specific distance from the earthquake. While it is correct to say that for each increase in 1 in the Richter Magnitude, there is a tenfold increase in amplitude of the wave, it is incorrect to say that each increase of 1 in Richter Magnitude represents a tenfold increase in the size of the Earthquake (as is commonly incorrectly stated by the Press). A better measure of the size of an earthquake is the amount of energy released by the earthquake. The amount of energy released is related to the Magnitude Scale by the following equation: Log E = 11.8 + 1.5 M Where Log refers to the logarithm to the base 10, E is the energy released in ergs, and M is the Magnitude. Anyone with a hand calculator can solve this equation by plugging in various values of M and solving for E, the energy released. I've done the calculation for you in the following table:

Magnitude Energy (ergs) Factor 1 2.0 x 1013 31 x 2 6.3 x 1014 3 2.0 x 1016 31 x 4 6.3 x 1017 5 2.0 x 1019 31 x 6 6.3 x 1020 7 2.0 x 1022 31 x 8 6.3 x 1023

From these calculations you can see that each increase in 1 in Magnitude represents a 31 fold increase in the amount of energy released. Thus, a magnitude 7 earthquake releases 31 times more energy than a magnitude 6 earthquake. A magnitude 8 earthquake releases 31 x 31 or 961 times more energy than a magnitude 6 earthquake. Although the Richter Magnitude is the scale most commonly reported when referring to the size of an earthquake, it has been found that for larger earthquakes a more accurate measurement of size is the moment magnitude, Mw.  The moment magnitude is a measure of the amount of strain energy released by the earthquake as determined by measurements of the shear strength of the rock and the area of the rupture surface that slipped during the earthquake. Note that it usually takes more than one seismographic station to calculate the magnitude of an earthquake. Thus you will hear initial estimates of earthquake magnitude immediately after an earthquake and a final assigned magnitude for the same earthquake that may differ from initial estimates, but is assigned after seismologists have had time to evaluate the data from numerous seismographic stations. The moment magnitude for large earthquakes is usually greater than the Richter magnitude for the same earthquake.  For example the Richter magnitude for the 1964 Alaska earthquake is usually reported as 8.6, whereas the moment magnitude for this earthquake is calculated at 9.2.  The largest earthquake ever recorded was in Chile in 1960 with a moment magnitude of 9.5, The Summatra earthquake of 2004 had a moment magnitude of 9.0.  Sometimes a magnitude is reported for an earthquake and no specification is given as to which magnitude (Richter or moment) is reported.  This obviously can cause confusion. But, within the last few years, the tendency has been to report the moment magnitude rather than the Richter magnitude.  The Hiroshima atomic bomb released an amount of energy equivalent to a moment magnitude 6 earthquake. Note that magnitude scales are open ended with no maximum or minimum. The largest earthquakes are probably limited by rock strength. Meteorite impacts could cause larger earthquakes than have ever been observed.

Frequency of Earthquakes of Different Magnitude Worldwide
Magnitude	Number of Earthquakes per Year	Description
> 8.5	0.3	Great
8.0 - 8.4	1
7.5 - 7.9	3	Major
7.0 - 7.4	15
6.6 - 6.9	56
6.0 - 6.5	210	Destructive
5.0 - 5.9	800	Damaging
4.0 - 4.9	6,200	Minor
3.0 - 3.9	49,000
2.0 - 2.9	300,000
0 - 1.9	700,000

Modified Mercalli Intensity Scale Note that the Richter magnitude scale results in one number for the size of the earthquake. Maximum ground shaking will occur only in the area of the epicenter of the earthquake, but the earthquake may be felt over a much larger area. The Modified Mercalli Scale was developed in the late 1800s  to assess the intensity of ground shaking and building damage over large areas. The scale is applied after the earthquake by conducting surveys of people's response to the intensity of ground shaking and destruction.

Intensity	Characteristic Effects	Richter Scale Equivalent
1	People do not feel any Earth movement	<3.4
2	A few people notice movement if at rest and/or on upper floors of tall buildings
3	People indoors feel movement. Hanging objects swing back and forth. People outdoors might not realize that an earthquake is occurring	4.2
4	People indoors feel movement. Hanging objects swing. Dishes, windows, and doors rattle. Feels like a heavy truck hitting  walls. Some people outdoors may feel movement. Parked cars rock.	4.3 - 4.8
5	Almost everyone feels movement. Sleeping people are awakened. Doors swing open/close. Dishes break.  Small objects move or are turned over. Trees shake. Liquids spill from open containers	4.9-5.4
6	Everyone feels movement. People have trouble walking. Objects fall from shelves. Pictures fall off walls. Furniture moves. Plaster in walls may crack. Trees and bushes shake. Damage slight in poorly built buildings.	5.5 - 6.1
7	People have difficulty standing. Drivers feel cars shaking. Furniture breaks. Loose bricks fall from buildings. Damage slight to moderate in well-built buildings; considerable in poorly built buildings.	5.5 - 6.1
8	Drivers have trouble steering. Houses not bolted down shift on foundations. Towers & chimneys twist and fall. Well-built buildings suffer slight damage. Poorly built structures severely damaged. Tree branches break. Hillsides crack if ground is wet. Water levels in wells change.	6.2 - 6.9
9	Well-built buildings suffer considerable damage. Houses not bolted down move off foundations. Some underground pipes broken. Ground cracks.   Serious damage to Reservoirs.	6.2 - 6.9
10	Most buildings & their foundations destroyed. Some bridges destroyed. Dams damaged. Large landslides occur. Water thrown on the banks of canals, rivers, lakes. Ground cracks in large areas. Railroad tracks bent slightly.	7.0 - 7.3
11	Most buildings collapse. Some bridges destroyed. Large cracks appear in the ground. Underground pipelines destroyed. Railroad tracks badly bent.	7.4 - 7.9
12	Almost everything is destroyed. Objects  thrown into the air. Ground moves in waves or ripples. Large amounts of rock may move.	>8.0

The Modified Mercalli Scale is shown in the table above. Note that correspondence between maximum intensity and Richter Scale magnitudeonly applies in the area around the epicenter. A given earthquake will have zones of different intensity all surrounding a zone of maximum intensity. The Mercalli Scale is very useful in examining the effects of an earthquake over a large area, because it will is responsive not only to the size of the earthquake as measured by the Richter scale for areas near the epicenter, but will also show the effects of the efficiency that seismic waves are transmitted through different types of material near the Earth's surface. The Mercalli Scale is also useful for determining the size of earthquakes that occurred before the modern seismographic network was available (before there were seismographic stations, it was not possible to assign a Magnitude).

What Happens During an Earthquake? Earthquakes produce several effects that cause damage and destruction. Some of these effects are the direct result of the ground shaking produced by the arrival of seismic waves and others are secondary effects.  Among these effects are the following:

Ground Shaking - Shaking of the ground caused by the passage of seismic waves near the epicenter of the earthquake is responsible for the collapse of most structures. The intensity of ground shaking depends on distance from the epicenter and on the type of bedrock underlying the area.   In general, loose unconsolidated sediment is subject to more intense shaking than solid bedrock. Damage to structures from shaking depends on the type of construction. Concrete and masonry structures, because they are brittle are more susceptible to damage than wood and steel structures, which are more flexible. Different kinds of shaking occur due to passage of different kinds of waves. As the P-waves arrive the ground will move up and down.  The S-waves produce waves that both move the ground up and down and back and forth in the direction of wave motion.  The Love waves shake the ground from side to side, and the Rayleigh waves create a rolling up and down motion (see figure 10.25 in your text).

Ground Rupture - Ground rupture only occurs along the fault zone that moves during the earthquake. Thus, structures that are built across fault zones may collapse, whereas structures built adjacent to, but not crossing the fault may survive. Fire - Fire is a secondary effect of earthquakes. Because power lines may be knocked down and because natural gas lines may rupture due to an earthquake, fires are often started closely following an earthquake. The problem is compounded if water lines are also broken during the earthquake since there will not be a supply of water to extinguish the fires once they have started. In the 1906 earthquake in San Francisco more than 90% of the damage to buildings was caused by fire. Landslides and Debris/Rock Falls - In mountainous regions subjected to earthquakes ground shaking may trigger rapid mass-wasting events like landslides, rock and debris falls, slumps, and debris avalanches.

Liquefaction -Liquefaction is a processes that occurs in water-saturated unconsolidated sediment due to shaking. In areas underlain by such material, the ground shaking causes the grains to loose grain to grain contact, and thus the material tends to flow.

You can demonstrate this process to yourself next time your go the beach. Stand on the sand just after an incoming wave has passed. The sand will easily support your weight and you will not sink very deeply into the sand if you stand still. But, if you start to shake your body while standing on this wet sand, you will notice that the sand begins to flow as a result of liquefaction, and your feet will sink deeper into the sand. Aftershocks - Earthquakes can change the stress state in rocks near the hypocenter and this may induce numerous earthquakes that occur after the main earthquake.  These are almost always smaller earthquakes, but they can be numerous and last for many months after the main earthquake.  Aftershocks are particularly dangerous because that can cause further damage to already damaged structures and make it unsafe for rescue efforts to be pursued. Tsunami - Tsunami are giant ocean waves that can rapidly travel across oceans. Earthquakes that occur along coastal areas can generate tsunami, which can cause damage thousands of kilometers away on the other side of the ocean. Tsunami can be generated by anything that disturbs a body of water.   This includes earthquakes that cause vertical offset of the sea floor, volcanic eruptions into a body of water, landslides into a body of water, underwater explosions, and meteorite impacts.  In general, the larger the earthquake, eruption, landslide, explosion or meteorite, the more likely it will be able to travel across an ocean.   Smaller events may, however cause a tsunami that affect areas in the vicinity of the triggering event.  Tsunami waves have wavelengths and velocities much higher that wind driven ocean waves.  Velocities are on the order of several hundred km/hr, similar to a jet airplane.  They usually are more than one wave, that hit the coastline tens of minutes to hours apart.  Although wave heights are barely perceptible in the open ocean, the waves become amplified as the approach the shore and may build to several tens of meters.  Thus, when the come ashore, the can flood areas far away from the coast.  Often the trough of a tsunami wave arrives before the crest,  This produces a phenomenon called drawdown where the ocean recedes from the normal shoreline by as much as a kilometer. Tsunami warning systems have been developed for the Pacific Ocean basin and, recently, the Indian Ocean where a tsunami killed over 250,000 people in 2004.  But, such warning systems depend on the ability to detect and forecast a tsunami after an earthquake occurs and may take several hours to come up with an accurate forecast of wave heights and travel time. Knowing something about these aspects of tsunami could save your life.  It suggests that If you are near the beach and feel an earthquake immediately get to higher ground.   Tsunami warnings require time and if you are near enough to the earthquake that generates a tsunami that you feel the earthquake, there may not be enough time for a warning to be sounded, nor will there be enough time to get out of the way once you see the wave approaching. If you are near the beach and see the ocean recede far offshore, immediately get to higher ground, as the receding ocean indicates that the trough of a tsunami wave has arrived and will be followed by the crest. If you survive the first wave of a tsunami, don't go back to the coast assuming the event is over.  Several waves are possible and any of them could be the largest of the waves.   Wait for authorities to issue an "all clear signal". Don't even consider "surfing the tsunami wave" or riding it out.  The waves are so powerful and last such a long time, that you would have little chance of surviving.

Where do Earthquakes Occur The distribution and frequency of earthquakes is referred to as seismicity. Most earthquakes occur along relatively narrow belts that coincide with plate boundaries (see figure 10.18 in your text).

This makes sense, since plate boundaries are zones along which lithospheric plates mover relative to one another. Earthquakes along these zones can be divided into shallow focus earthquakes that have focal depths less than about 70 km and deep focus earthquakes that have focal depths between 75 and 700 km.

Earthquakes at Diverging Plate Boundaries Diverging plate boundaries are zones where two plates move away from each other, such as at oceanic ridges. In such areas the lithosphere is in a state of tensional stress and thus normal faults and rift valleys occur. Earthquakes that occur along such boundaries show normal fault motion and tend to be shallow focus earthquakes, with focal depths less than about 20 km. Such shallow focal depths indicate that the brittle lithosphere must be relatively thin along these diverging plate boundaries. Earthquakes at Converging Plate Boundaries - Convergent plate boundaries are boundaries where two plates run into each other. Thus, they tend to be zones where compressional stresses are active and thus reverse faults or thrust faults are common. There are two types of converging plate boundaries. (1) subduction boundaries, where oceanic lithosphere is pushed beneath either oceanic or continental lithosphere; and (2) collision boundaries where two plates with continental lithosphere collide. Subduction boundaries -At subduction boundaries cold oceanic lithosphere is pushed back down into the mantle where two plates converge at an oceanic trench. Because the subducted lithosphere is cold, it remains brittle as it descends and thus can fracture under the compressional stress. When it fractures, it generates earthquakes that define a zone of earthquakes with increasing focal depths beneath the overriding plate. This zone of earthquakes is called the Benioff Zone. Focal depths of earthquakes in the Benioff Zone can reach down to 700 km.

Collision boundaries - At collisional boundaries two plates of continental lithosphere collide resulting in fold-thrust mountain belts. Earthquakes occur due to the thrust faulting and range in depth from shallow to about 200 km. Earthquakes at Transform Fault Boundaries Transform fault boundaries are plate boundaries where lithospheric plates slide past one another in a horizontal fashion. The San Andreas Fault of California is one of the longer transform fault boundaries known. Earthquakes along these boundaries show strike-slip motion on the faults and tend to be shallow focus earthquakes with depths usually less than about 50 km. Intraplate Earthquakes - These are earthquakes that occur in the stable portions of continents that are not near plate boundaries.  Many of them occur as a result of re-activation of ancient faults, although the causes of some intraplate earthquakes are not well understood. Examples - New Madrid Region, Central U.S., Charleston South Carolina, Along St. Lawrence River - U.S. - Canada Border.

Earthquake RiskThe risk that an earthquake will occur close to where you live depends on whether or not tectonic activity that causes deformation is occurring within the crust of that area. For the U.S., the risk is greatest in the most tectonically active area, that is near the plate margin in the Western U.S. Here, the San Andreas Fault which forms the margin between the Pacific Plate and the North American Plate, is responsible for about 1 magnitude 8 or greater earthquake per century. Also in the western U.S. is the Basin and Range Province where extensional stresses in the crust have created many normal faults that are still active. Historically, large earthquakes have also occurred in the area of New Madrid, Missouri;and Charleston, South Carolina. (See figure 10.34 in your text). Why earthquakes occur in these other areas is not well understood. If earthquakes have occurred before, they are expected to occur again.

Long-Term Forecasting Long-term forecasting is based mainly on the knowledge of when and where earthquakes have occurred in the past.  Thus, knowledge of present tectonic setting, historical records, and geological records are studied to determine locations and recurrence intervals of earthquakes.  Two methods of earthquake forecasting are being employed - paleoseismology and seismic gaps. Paleoseismology - the study of prehistoric earthquakes.  Through study of the offsets in sedimentary layers near fault zones, it is often possible to determine recurrence intervals of major earthquakes prior to historical records.  If it is determined that earthquakes have recurrence intervals of say 1 every 100 years, and there are no records of earthquakes in the last 100 years, then a long-term forecast can be made and efforts can be undertaken to reduce seismic risk. Seismic gaps - A seismic gap is a zone along a tectonically active area where no earthquakes have occurred recently, but it is known that elastic strain is building in the rocks.  If  a seismic gap can be identified, then it might be an area expected to have a large earthquake in the near future.

Short-Term Prediction Short-term predication involves monitoring of processes that occur in the vicinity of earthquake prone faults for activity that signify a coming earthquake.   Anomalous events or processes that may precede an earthquake are calledprecursor events and might signal a coming earthquake. Despite the array of possible precursor events that are possible to monitor, successful short-term earthquake prediction has so far been difficult to obtain.   This is likely because: the processes that cause earthquakes occur deep beneath the surface and are difficult to monitor. earthquakes in different regions or along different faults all behave differently, thus no consistent patterns have so far been recognized Among the precursor events that may be important are the following: Ground Uplift and Tilting  - Measurements taken in the vicinity of active faults sometimes show that prior to an earthquake the ground is uplifted or tilts due to the swelling of rocks caused by strain building on the fault.  This may lead to the formation of numerous small cracks (called microcracks).  This cracking in the rocks may lead to small earthquakes called foreshocks. Foreshocks - Prior to a 1975 earthquake in China, the observation of numerous foreshocks led to successful prediction of an earthquake and evacuation of the city of the Haicheng.  The magnitude 7.3 earthquake that occurred, destroyed half of the city of about 100 million inhabitants, but resulted in only a few hundred deaths because of the successful evacuation.  Water Level in Wells - As rocks become strained in the vicinity of a fault, changes in pressure of the groundwater (water existing in the pore spaces and fractures in rocks) occur.  This may force the groundwater to move to higher or lower elevations, causing changes in the water levels in wells. Emission of Radon Gas - Radon is an inert gas that is produced by the radioactive decay of uranium and other elements in rocks.  Because Radon is inert, it does not combine with other elements to form compounds, and thus remains in a crystal structure until some event forces it out.  Deformation resulting from strain may force the Radon out and lead to emissions of Radon that show up in well water.  The newly formed microcracks discussed above could serve as pathways for the Radon to escape into groundwater.  Increases in the amount of radon emissions have been reported prior to some earthquakes Strange Animal Behavior - Prior to a magnitude 7.4 earthquake in Tanjin, China, zookeepers reported unusual animal behavior.  Snakes refusing to go into their holes, swans refusing to go near water, pandas screaming, etc.  This was the first systematic study of this phenomenon prior to an earthquake.  Although other attempts have been made to repeat a prediction based on animal behavior, there have been no other successful predictions.

Controlling Earthquakes Although no attempts have yet been made to control earthquakes, earthquakes have been known to be induced by human interaction with the Earth.  This suggests that in the future earthquake control may be possible. Examples of human induced earthquakes For ten years after construction of the Hoover Dam in Nevada blocking the Colorado River to produce Lake Mead, over 600 earthquakes occurred, one with magnitude of 5 and 2 with magnitudes of 4. In the late 1960s toxic waste injected into hazardous waste disposal wells at Rocky Flats, near Denver apparently caused earthquakes to occur in a previously earthquake quiet area.  The focal depths of the quakes ranged between 4 and 8 km, just below the 3.8 km-deep wells. Nuclear testing in Nevada set off thousands of aftershocks after the explosion of a 6.3 magnitude equivalent underground nuclear test.  The largest aftershocks were about magnitude 5.  In the first two examples the increased seismicity was apparently due to increasing fluid pressure in the rocks which resulted in re-activating older faults by increasing strain. The problem, however, is that of the energy involved.  Remember that for every increase in earthquake magnitude there is about a 30 fold increase in the amount of energy released.  Thus, in order to release the same amount of energy as a magnitude 8 earthquake, 30 magnitude 7 earthquakes would be required.  Since magnitude 7 earthquakes are still very destructive, we might consider generating smaller earthquakes.   If we say that a magnitude 4 earthquake might be acceptable, how many magnitude 4 earthquakes are required to release the same amount of energy as a magnitude 8 earthquake?   Answer 30 x 30 x 30 x 30 =810,000!  Still, in the future it may be possible to control earthquakes either with explosions to gradually reduce the stress or by pumping fluids into the ground.

Mitigating for Earthquake Hazards Many seismologists have said that "earthquakes don't kill people, buildings do". This is because most deaths from earthquakes are caused by buildings or other human construction falling down during an earthquake. Earthquakes located in isolated areas far from human population rarely cause any deaths. Thus, in earthquake prone areas like California, there are strict building codes requiring the design and construction of buildings and other structures that will withstand a large earthquake. While this program is not always completely successful, one fact stands out to prove its effectiveness. In 1986 an earthquake near San Francisco, California with a Richter Magnitude of 7.1 killed about 40 people. Most were killed when a double decked freeway collapsed. About 10 months later, an earthquake with magnitude 6.9 occurred in the Armenia, where no earthquake proof building codes existed. The death toll in the latter earthquake was about 25,000! Another contrast occurred in 2010. On January 12, an earthquake of Moment Magnitude 7.0 occurred in Haiti.  The country is one of the poorest on earth, had no earthquake resistant building codes, and most of the construction was poorly reinforced concrete.   The destruction was massive with an estimated 250,000 deaths.  On February 27, a Moment Magnitude 8.8 earthquake occurred in Chile, a country where earthquake resistant building codes were enforced.  The death toll from this larger earthquake was about 520, again, proving the effectiveness of building codes.

How Seismic Waves Help Understand Earth's Internal Structure Much of what we know about the interior of the Earth comes from knowledge of seismic wave velocities and their variation with depth in the Earth. Recall that body wave velocities are as follows: Where K = incompressibility μ = rigidity ρ = densityIf the properties of the earth, i.e. K,μ, andρwhere the same throughout, then Vp and Vs would be constant throughout the Earth and seismic waves would travel along straight line paths through the Earth. We know however that density must change with depth in the Earth, because the density of the Earth is 5,200 kg/cubic meter and density of crustal rocks is about 2,500 kg/cubic meter. If the density were the only property to change, then we could make estimates of the density, and predict the arrival times or velocities of seismic waves at any point away from an earthquake. Observations do not follow the predictions, so, something else must be happening. In fact we know that K, μ, and ρ change due to changing temperatures, pressures and compositions of material. The job of seismology is, therefore, to use the observed seismic wave velocities to determine how K,μ, and ρ change with depth in the Earth, and then infer how pressure , temperature, and composition change with depth in the Earth. In other words to tell us something about the internal structure of the Earth. Reflection and Refraction of Seismic Waves If composition (or physical properties) change abruptly at some interface, then seismic wave will both reflect off the interface and refract (or bend) as they pass through the interface. Two cases of wave refraction can be recognized. If the seismic wave velocity in the rock above an interface is less than the seismic wave velocity in the rock below the interface, the waves will be refracted or bent upward relative to their original path.

If the seismic wave velocity decreases when passing into the rock below the interface, the waves will be refracted down relative to their original path.

If the seismic wave velocities gradually increase with depth in the Earth, the waves will continually be refracted along curved paths that curve back toward the Earth's surface.

One of the earliest discoveries of seismology was a discontinuity at a depth of 2900 km where the velocity of P-waves suddenly decreases. This boundary is the boundary between the mantle and the core and was discovered because of a zone on the opposite side of the Earth from an earthquake focus receives no direct P-waves because the P-waves are refracted inward as a result of the sudden decrease in velocity at the boundary.
This zone is called a P-wave shadow zone.
This discovery was followed by the discovery of an S-wave shadow zone. The S-wave shadow zone occurs because no S-waves reach the area on the opposite side of the Earth from the focus. Since no direct S-waves arrive in this zone, it implies that no S-waves pass through the core. This further implies the velocity of S-wave in the core is 0. In liquids μ = 0, so S-wave velocity is also equal to 0. From this it is deduced that the core, or at least part of the core is in the liquid state, since no S-waves are transmitted through liquids.

Thus, the S-wave shadow zone is best explained by a liquid outer core.

Seismic Wave Velocities in the Earth Over the years seismologists have collected data on how seismic wave velocities vary with depth in the Earth. Distinct boundaries, called discontinuities are observed when there is sudden change in physical properties or chemical composition of the Earth. From these discontinuities, we can deduce something about the nature of the various layers in the Earth. As we discussed way back at the beginning of the course, we can look at the Earth in terms of layers of differing chemical composition, and layers of differing physical properties.

Layers of Differing Composition - The Crust - Mohorovicic discovered boundary the boundary between crust and mantle, thus it is named the Mohorovicic Discontinuity or Moho, for short. The composition of the crust can be determined from seismic waves by comparing seismic wave velocities measured on rocks in the laboratory with seismic wave velocities observed in the crust. Then from travel times of waves on many earthquakes and from many seismic stations, the thickness and composition of the crust can be inferred. In the ocean basins crust is about 8 to 10 km thick, and has a composition that is basaltic. Continental crust varies between 20 and 60 km thick. The thickest continental crust occurs beneath mountain ranges and the thinnest beneath lowlands. The composition of continental crust varies from granitic near the top to gabbroic near the Moho. The Mantle - Seismic wave velocities increase abruptly at the Moho. In the mantle wave velocities are consistent with a rock composition of peridotite which consists of olivine, pyroxene, and garnet. . The Core - At a depth of 2900 Km P-wave velocities suddenly decrease and S-wave velocities go to zero. This is the top of the outer core. As discussed above, the outer core must be liquid since S-wave velocities are 0. At a depth of about 4800 km the sudden increase in P-wave velocities indicate a solid inner core. The core appears to have a composition consistent with mostly Iron with small amounts of Nickel.

Layers of Different Physical Properties At a depth of about 100 km there is a sudden decrease in both P and S-wave velocities. This boundary marks the base of the lithosphere and the top of the asthenosphere. The lithosphere is composed of both crust and part of the upper mantle. It is a brittle layer that makes up the plates in plate tectonics, and appears to float and move around on top of the more ductile asthenosphere. At the top of the asthenosphere is a zone where both P- and S-wave velocities are low. This zone is called the Low-Velocity Zone (LVZ). It is thought that the low velocities of seismic waves in this zone are caused by temperatures approaching the partial melting temperature of the mantle, causing the mantle in this zone to behave in a very ductile manner. At a depth of 400 km there is an abrupt increase in the velocities of seismic waves, thus this boundary is known as the 400 - Km Discontinuity. Experiments on mantle rocks indicate that this represents a temperature and pressure where there is a polymorphic phase transition, involving a change in the crystal structure of Olivine, one of the most abundant minerals in the mantle. Another abrupt increase in seismic wave velocities occurs at a depth of 670 km. It is uncertain whether this discontinuity, known as the 670 Km Discontinuity, is the result of a polymorphic phase transition involving other mantle minerals or a compositional change in the mantle, or both.  Seismic Tomography Most of you are aware of the techniques used in modern medicine to see inside the human body.   These are things like CT scans, ultrasound, and X-rays.  All them use waves, either sound waves or electromagnetic waves, that penetrate the body and reflect and refract from and through body parts that have different physical properties.   The techniques require a source of waves with enough energy to penetrate, the ability to generate these waves continuously in places that will penetrate the area of interest, and the ability to detect the resulting reflected and refracted waves when they emerge.  Similar imaging can be done for the earth, but it is much more complicated. Seismic waves from a large earthquake can penetrate the earth, but each earthquake is a single point source for the waves.   Seismometers can detect the waves when they emerge, but seismometers are not placed everywhere on the earth's surface.   Nevertheless, if data is collected over many years, the information can be used to produce an image of the interior of the earth.   Such images are sill pretty primitive, but allow us to see areas that are hotter than their surroundings, where seismic wave velocities are slower and areas that are cooler than their surroundings where velocities are higher.

Examples of questions on this material that could be asked on an exam Define the following terms (a), stress (b) confining stress, (c) differential stress, (d) tensional stress (e) compressional stress, (e) strain (f) liquifaction, (g) fault creep, (h) Benioff Zone. What are the three stages of deformation that all materials go through as stress is increased? What is the difference between a brittle material and a ductile material? Explain the following types of faults: (a) normal fault, (b) reverse fault, (c) thrust fault, (d) strike-slip fault, and (e) transform fault. Explain the elastic rebound theory on the cause earthquakes. What is the difference between the epicenter and the focus of an earthquake. What are seismic waves and what is the difference between a P-wave, an S-wave and a Surface waves? For each increase of magnitude by a factor of 1, how much more energy is released? What is the difference between Richter magnitude and Moment magnitude and which of these scales is a more accurate measure of the energy released by large earthquakes?

Pleochroism

Pleochroism is defined as the change in colour of a mineral, in plane light, on rotating the stage. It occurs when the wavelengths of the ordinary & extraordinary rays are absorbed differently on passing through a mineral, resulting in different wavelengths of light passing the mineral.
Coloured minerals, whether uniaxial or biaxial, are generally pleochroic.
To describe the pleochroism for uniaxial minerals must specify the colour which corresponds to the ordinary and extraordinary rays.

• e.g. Tourmaline, Hexagonal mineral
  • omega = dark green
  • epsilon = pale green

If the colour change is quite distinct the pleochroism is said to be strong.
If the colour change is minor = weak pleochroism.
For coloured uniaxial minerals, sections cut perpendicular to the c axis will show a single colour, corresponding to ordinary ray.
Sections parallel to the c crystallographic axis will exhibit the widest colour variation as both omega and epsilon are present.see pleochroism in the earlier published uniaxial and biaxial minerals

Extinction in minerals

We wish to examine other properties of minerals which are useful in the identification of unknown minerals.

Anisotropic minerals go extinct between crossed polars every 90° of rotation. Extinction occurs when one vibration direction of a mineral is parallel with the lower polarizer. As a result no component of the incident light can be resolved into the vibration direction of the upper polarizer, so all the light which passes through the mineral is absorbed at the upper polarizer, and the mineral is black.

Upon rotating the stage to the 45° position, a maximum component of both the slow and fast ray is available to be resolved into the vibration direction of the upper polarizer. Allowing a maximum amount of light to pass and the mineral appears brightest.

The only change in the interference coloursis that they get brighter or dimmer withrotation, the actual colours do not change.

Many minerals generally form elongate grains and have an easily recognizable cleavage direction, e.g. biotite, hornblende, plagioclase.

The extinction angle is the angle between the length or cleavage of a mineral and the minerals vibration directions.

The extinction angles when measured on several grains of the same mineral, in the same thin section, will be variable. The angle varies because of the orientation of the grains. The maximum extinction angle recorded is diagnostic for the mineral.

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Types of Extinction

1. Parallel ExtinctionThe mineral grain is extinct when the cleavage or length is aligned with one of the crosshairs.
   The extinction angle (EA) = 0°
   

e.g.

• orthopyroxene
• biotite

2. Inclined ExtinctionThe mineral is extinct when the cleavage is at an angle to the crosshairs.
   EA > 0°

e.g.

• clinopyroxene
• hornblende

3. Symmetrical ExtinctionThe mineral grain displays two cleavages or two distinct crystal faces. It is possible to measure two extinction angles between each cleavage or face and the vibration directions. If the two angles are equal then Symmetrical extinction exists.
   EA₁ = EA₂
   

e.g.

• amphibole
• calcite

4. No CleavageMinerals which are not elongated or do not exhibit a prominent cleavage will still go extinct every 90° of rotation, but there is no cleavage or elongation direction from which to measure the extinction angle.
   

e.g.

• quartz
• olivine

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Exceptions to Normal ExtinctionPatterns

Different portions of the same grain may go extinct at different times, i.e. they have different extinction angles. This may be caused by chemical zonation or strain.

Chemical zonation

The optical properties of a mineral vary with the chemical composition resulting in varying extinction directions for a mineral. Such minerals are said to be zoned.
e.g. plagioclase, olivine

Strain

During deformation some grains become bent, resulting in different portions of the same grain having different orientations, therefore they go extinct at different times.
e.g. quartz, plagioclase

interference colour

                  INTERFERENCE COLOUR 

An interference colour is the colour observed for a mineral in thin-section under crosed polarized light. Interference colours can provide important clues to the identity of a mineral or even the orientation of the crystal in the section. Interference colours are produced because plane polarised light passing through a mineral section splits into two rays vibrating at 90 degrees to each other. The two rays travel at a different velocities depending on the refractive index of a mineral in the direction of the vibration. The rays are, therefore, called the fast ray and the slow ray. Because the fast and slow ray travel at different velocities they become out of phase and when they recombine at the analyser they interfere to produce light of a characteristic wavelength. Interference colours change with the orientation of the mineral section for anisotropic minerals because the refractive indices of the mineral are different in different directions. Isotropic minerals do not show interference colours (they are black in cross polarised light) since their refractive index is the same in every direction. 
 
 The interference colour of a section of a mineral can be predicted from the birefringence- the difference between the refractive indices in the vibration directions of the fast and slow ray. Usually optical mineralogy text books provide the range of principal refractive indices for the principal axes of the mineral, and the orientation of these axes relative to the crystallographic axes of the mineral. The interference colour produced also changes with the thickness of the section.

sedimentary rocks

Sedimentary Rocks Rivers, oceans, winds, and rain runoff all have the ability to carry the particles washed off of eroding rocks. Such material, called detritus, consists of fragments of rocks and minerals. When the energy of the transporting current is not strong enough to carry these particles, the particles drop out in the process of sedimentation. This type of sedimentary deposition is referred to asclastic sedimentation. Another type of sedimentary deposition occurs when material is dissolved in water, and chemically precipitates from the water. This type of sedimentation is referred to aschemical sedimentation. A third process can occur, wherein living organisms extract  ions dissolved in water to make such things as shells and bones.  This type of sedimentation is called biochemical sedimentation.  The accumulation of plant matter, such as at the bottom of a swamp, is referred to as organic sedimentation.  Thus, there are 4 major types of sedimentary rocks: Clastic Sedimentary Rocks, Chemical Sedimentary Rocks,Biochemical Sedimentary Rocks, andOrganic Sedimentary Rocks. Clastic Sediments and Sedimentary Rocks

The formation of a clastic sediment and sedimentary rocks involves five processes: Weathering - The first step is transforming solid rock into smaller fragments or dissolved ions by physical and chemical weathering as discussed in the last lecture. Erosion - Erosion is actually many process which act together to lower the surface of the earth.  In terms of producing sediment, erosion begins the transpiration process by moving the weathered products from their original location.  This can take place by gravity (mass wasting events like landslides or rock falls), by running water. by wind, or by moving ice.  Erosion overlaps with transpiration. Transportation -  Sediment can be transported by sliding down slopes, being picked up by the wind, or by being carried by running water in streams, rivers, or ocean currents. The distance the sediment is transported and the energy of the transporting medium all leave clues in the final sediment that tell us something about the mode of transportation.

Deposition -  Sediment is deposited when the energy of the transporting medium becomes too low to continue the transport process. In other words, if the velocity of the transporting medium becomes too low to transport sediment, the sediment will fall out and become deposited. The final sediment thus reflects the energy of the transporting medium.  Lithification (Diagenesis) - Lithification is the process that turns sediment into rock. The first stage of the process is compaction. Compaction occurs as the weight of the overlying material increases. Compaction forces the grains closer together, reducing pore space and eliminating some of the contained water. Some of this water may carry mineral components in solution, and these constituents may later precipitate as new minerals in the pore spaces. This causes cementation, which will then start to bind the individual particles together. Classification - Clastic sedimentary particles and sedimentary rocks are classified in terms of grain size and shape, among other factors.

Name of Particle	Size Range	Loose Sediment	Consolidated Rock
Boulder	>256 mm	Gravel	Conglomerate or Breccia (depends on rounding)
Cobble	64 - 256 mm	Gravel
Pebble	2 - 64 mm	Gravel
Sand	1/16 - 2mm	Sand	Sandstone
Silt	1/256 - 1/16 mm	Silt	Siltstone
Clay	<1/256 mm	Clay	Claystone, mudstone, and shale

In general, the coarser sediment gets left behind by the transportation process. Thus, coarse sediment is usually found closer to its source and fine grained sediment is found farther from the source. Textures of Clastic Sedimentary Rocks When sediment is transported and deposited, it leaves clues to the mode of transport and deposition. For example, if the mode of transport is by sliding down a slope, the deposits that result are generally chaotic in nature, and show a wide variety of particle sizes. Grain size and the interrelationship between grains gives the resulting sediment texture. Thus, we can use the texture of the resulting deposits to give us clues to the mode of transport and deposition.

Sorting - The degree of uniformity of grain size. Particles become sorted on the basis of density, because of  the energy of the transporting medium.  High energy currents can carry larger fragments.  As the energy decreases, heavier particles are deposited and lighter fragments continue to be transported.  This results in sorting due to density.

If the particles have the same density, then the heavier particles will also be larger, so the sorting will take place on the basis of size.  We can classify this size sorting on a relative basis -  well sorted to poorly sorted. Sorting gives clues to the energy conditions of the transporting medium from which the sediment was deposited.
Examples Beach deposits and wind blown deposits generally show good sorting because the energy of the transporting medium is usually constant. Stream  deposits are usually poorly sorted because the energy (velocity) in a stream varies with position in the stream and time.

Rounding - During the transportation process, grains may be reduced in size due to abrasion.  Random abrasion results in the eventual rounding off of the sharp corners and edges of grains.  Thus, rounding of grains gives us clues to the amount of time a sediment has been in the transportation cycle.  Rounding is classified on relative terms as well.

Sediment Maturity Sediment Maturity refers to the length of time that the sediment has been in the sedimentary cycle.   Texturally mature sediment is sediment that is well rounded, (as rounding increases with transport distance and time) and well sorted (as sorting gets better as larger clasts are left behind and smaller clasts are carried away.  Because the weathering processes continues during sediment transport, mineral grains that are unstable near the surface become less common as the distance of transport or time in the cycle increases.  Thus compositionally mature sediment is composed of only the most stable minerals. For example a poorly sediment containing glassy angular volcanic fragments, olivine crystals and plagioclase is texturally immature because the fragments are angular, indicating they have not been transported very far and the sediment is poorly sorted, indicating that little time has been involved in separating larger fragments from smaller fragments. It is compositionally immature because it contains unstable glass along with minerals that are not very stable near the surface - olivine and plagioclase. On the other hand a well sorted beach sand consisting mainly of well rounded quartz grains is texturally mature because the grains are rounded, indicating a long time in the transportation cycle, and the sediment is well sorted, also indicative of the long time required to separate the coarser grained material and finer grained material from the sand.  The beach sand is compositionally mature because it is made up only of quartz which is very stable at the earth's surface. Types of Clastic Sedimentary Rocks We next look at various clastic sedimentary rocks that result from lithification of sediment. Conglomerates and Breccias Conglomerate and Breccia are rocks that contain an abundance of coarse grained clasts (pebbles, cobbles, or boulders). In a conglomerate, the coarse grained clasts are well rounded, indicating that they spent considerable time in the transportation process and were ultimately deposited in a high energy environment capable of carrying the large clasts. In a breccia, the coarse grained clasts are very angular, indicating the the clasts spent little time in the transportation cycle. Sandstones A Sandstone is made of sand-sized particles and forms in many different depositional settings. Texture and composition permit historic interpretation of the transport and depositional cycle and sometimes allows determination of the source. Quartz is, by far, the dominant mineral in sandstones. Still there are other varieties.  A Quartz arenite – is nearly 100% quartz grains.   An Arkose contains abundant feldspar.  In a lithic sandstone, the grains are mostly small rock fragments.  A Wacke is a sandstone that contains more than 15% mud (silt and clay sized grains).. Sandstones are one of the most common types of sedimentary rocks. Mudrocks Mudrocks are made of fine grained clasts (silt and clay sized) .  A siltstone is one variety that consists of silt-sized fragments.  A shale is composed of clay sized particles and is a rock that tends to break into thin flat fragments (See figure 7.4e in your text).  A mudstone is similar to a shale, but does not break into thin flat fragments. Organic-rich shales are the source of petroleum. Fine grained clastics are deposited in non-agitated water, calm water, where there is little energy to continue to transport the small grains.   Thus mudrocks form in deep water ocean basins and lakes.

Biochemical and Organic Sediments and Sedimentary Rocks Biochemical and Organic sediments and sedimentary rocks are those derived from living organisms.  When the organism dies, the remains can accumulate to become sediment or sedimentary rock.  Among the types of rock produced by this process are: Biochemical Limestone - calcite (CaCO3) is precipitated by organisms usually to form a shell or other skeletal structure.  Accumulation of these skeletal remains results in a limestone. Sometimes the fossilized remains of the organism are preserved in the rock, other times recrystallization during lithification has destroyed the remains. Limestones are very common sedimentary rocks. Biochemical Chert - Tiny silica secreting planktonic organism like Radiolaria and Diatoms can accumulate on the sea floor and recrystallize during lithification to form biochemical chert.  The recrystallization results in a hard rock that is usually seen as thin beds (see figure 7.22a in your test). Diatomite - When diatoms accumulate and do not undergo recrystallization, they form a white rock called diatomite as seen if the White Cliffs of Dover (see figure 7.20b in your text). Coal - Coal is an organic rock made from organic carbon that is the remains of fossil plant matter. It accumulates in lush tropical wetland settings and requires deposition in absence of Oxygen. It is high in carbon and can easily be burned to obtain energy. Chemical Sediments and Sedimentary Rocks Dissolved ions released into water by the weathering process are carried in streams or groundwater. Eventually these dissolved ions end in up in the ocean, explaining why sea water is salty.  When water evaporates or the concentration of the ions get too high as a result of some other process, the ions recombine by chemical precipitation to form minerals that can accumulate to become chemical sediments and chemical sedimentary rocks.  Among these are: Evaporites - formed by evaporation of sea water or lake water.  Produces halite (salt) and gypsum deposits by chemical precipitation as concentration of solids increases due to water loss by evaporation.  This can occur in lakes that have no outlets (like the Great Salt Lake) or restricted ocean basins, like has happened in the Mediterranean Sea or the Gulf of Mexico in the past. Travertine  - Groundwater containing dissolve Calcium and bicarbonate ions can precipitate calcite to form a chemically precipitated limestone, called travertine. This can occur in lakes, hot springs, and caves. Dolostones - Limestone that have been chemically modified by Mg-rich fluids flowing through the rock are converted to dolostones. CaCO3 is recrystallized to a new mineral dolomite CaMg(CO3)2. Chemical Cherts - Groundwater flowing through rock can precipitate SiO2 to replace minerals that were present.   This produces a non-biogenic chert.   There are many varsities of such chert that are given different names depending on their attributes,  For example: Flint – Black or gray from organic matter. Jasper – Red or yellow from Fe oxides. Petrified wood – Wood grain preserved by silica. Agate – Concentrically layered rings

Sedimentary Structures As mentioned previously, all stages of the sedimentary cycle leave clues to processes that were operating in the past.   Perhaps the most easily observable clues are structures left by the depositional process.   We here discuss sedimentary structures and the information that can be obtained from these structures.   Stratification and Bedding Because sediment is deposited in low lying areas that often extend over wide areas, successive depositional events produce layers called bedding or stratification that is usually the most evident feature of sedimentary rocks. The layering can be due to differences in color of the material, differences in grain size, or differences in mineral content or chemical composition. All of these differences can be related to differences in the environment present during the depositional events. (see figure 7.10 in your text). A series of beds are referred to as strata. A sequence of strata that is sufficiently unique to be recognized on a regional scale is termed a formation. A formation is the fundamental geologic mapping unit.  (See figure 7.11 in your text). Rhythmic Layering - Alternating parallel layers having different properties. Sometimes caused by seasonal changes in deposition (Varves). i.e. lake deposits wherein coarse sediment is deposited in summer months and fine sediment is deposited in the winter when the surface of the lake is frozen.

Cross Bedding - Sets of beds that are inclined relative to one another.  The beds are inclined in the direction that the wind or water was moving at the time of deposition.  Boundaries between sets of cross beds usually represent an erosional surface. Very common in beach deposits, sand dunes, and river deposited sediment.

Graded Bedding  - As current velocity decreases, first the larger or more dense particles are deposited followed by smaller particles.  This results in bedding showing a decrease in grain size from the bottom of the bed to the top of the bed.  Sediment added as a pulse of turbid water. As pulse wanes, water loses velocity and sediments settle. Coarsest material settles first, medium next, then fine. Multiple graded-bed sequences called turbidites (see figure 7.14 in your text).

Non-sorted Sediment - Sediment showing a mixture of grain sizes results from such things as rockfalls, debris flows, mudflows, and deposition from melting ice.

Ripple Marks  - Water flowing over loose sediment creates bedforms by moving sediment with the flow.

Bedforms are linked to flow velocity and sediment size. Ripples are characteristic of shallow water deposition and can also be caused by wind.  blowing over the surface.   Sand dunes are similar, but on a larger scale.  Ripples are commonly preserved in sedimentary rocks. Asymmetric ripples (as shown above) indicate flow direction,with the steep slope on the down - current direction.   Ripples persevered in ancient rocks can also be indicators of up/down direction in the original sediment. Symmetric ripples form as a result of constant wave energy oscillating back and forth.

Mudcracks - result from the drying out of wet sediment at the surface of the Earth.  The cracks form due to shrinkage of the sediment as it dries. When present in rock, they indicate that the surface was exposed at the earth's surface and then rapidly buried.

Sole Marks  - Flutes are troughs eroded in soft sediment that can become filled with mud. Both the flutes and the resulting casts (called flute casts) can be preserved in rock. Raindrop Marks - pits (or tiny craters) created by falling rain. If present, this suggests that the sediment was exposed to the surface of the Earth just prior to burial. Fossils - Remains of once living organisms.  Probably the most important indicator of the environment of deposition. Different species usually inhabit specific environments. Because life has evolved - fossils give clues to relative age of the sediment. Can also be important indicators of past climates. Rock Color Sulfides along with buried organic matter give rocks a dark color.  Indicates deposition in a reducing environment. Deposition in oxidizing environment produces red colored iron oxides and is often indicative of deposition in a non-marine environment.  Such red colored rocks are often referred to as red beds.

Sedimentary Environments If we look at various environments now present on Earth, we can find characteristics in the sediment that are unique to each environment. If we find those same characteristics in sedimentary rocks, it allows us to interpret the environment of the past.  Each environment has its own energy regime and sediment delivery, transport and depositional conditions that are reflected in the sediment deposited. Sedimentary Environments can be divided into the following

Terrestrial (Non-marine) environments Glacial Alluvial fans Sand Dunes Mountain Streams Lakes Rivers Marine environments Deltas Coastal Beaches Shallow Marine Clastics Shallow Marine Carbonates Deep Marine We will cover most of these environments in more detail later in the course. For now familiarize yourself with each of these by reading pages 202 to 206 in your text. Transgressions and Regressions Throughout geologic history sea level has risen and fallen by as much as a few hundred meters many times.  These changes are the result of changes earth's climate or changes in the shape of the sea floor as a result of tectonics. When sea level rises, the coast migrates inland. This is called a Transgression.  Beach sand gets buried by marine sediments and the sea floor subsides due to the weight of the sediment. During a transgression, the beach sand forms an extensive layer, but does not all have the same age. When sea level falls, the coast migrates seaward. This is called aRegression. The sedimentary sequence then repeats itself in a vertical sense as the sedimentary environment migrates back and forth. See figure 7.21 in your text. Diagenesis LIthification of sediment into sedimentary rocks takes place after the sediment has been deposited and buried.   The processes by which the sediment becomes lithified into a hard sedimentary rock is called diagenesis and includes all physical, chemical and biological processes that act on the sediment.   The first step in diagenesis is the compaction of the sediment and loss of water as a result of the weight of the overlying sediment.  Compaction and burial may cause recrystallization of the minerals to make the rock even harder. Fluids flowing through the rock and organisms may precipitate new minerals in the pore spaces between grains to form a cement that holds the sediment together.   Common cements include quartz, calcite, and hematite. Other conditions present during diagenesis, such as the presence of absence of free oxygen may cause other alterations to the original sediment. In an environment where there is excess oxygen (Oxidizing Environment) organic remains will be converted to carbon dioxide and water. Iron will change from Fe2+ to Fe3+, and will change the color of the sediment to a deep red (rust) color. In an environment where there is a depletion of oxygen (Reducing Environment), organic material may be transformed to solid carbon in the form of coal, or may be converted to hydrocarbons, the source of petroleum. Diagenesis is also a response to increasing the temperature and pressure as sediment gets buried deeper.  As temperature increases beyond about 200oC, we enter the realm of metamorphism