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Western University
Geography 2240A/B
Philip Egberts

Chapter 4 Earthquakes: The Basics 1.0 Introduction An accounting of the worst natural disasters in historical time places earthquakes and their consequences well within the 10 deadliest of processes. For example, in 1201 a devastating earthquake roared through Syria and Egypt killing approximately 1.1 million people. The population density in 1201 was rather dramatically less than now, so this was certainly an awesome catastrophe. The worst earthquake in China – which experiences many – seems to have been in 1556 when about 830,000 were killed in Shansi Province. Almost every month we read of someone killed by falling structures loosened by an earthquake. Of course, the deadly tsunami of December 2004 – with a dead + missing and presumed dead toll of about 250,000 – was instigated by an enormous earthquake. The earthquake in Haiti in 2010 was truly a disaster; had the country not been so poor, had infrastructure not been so weakly constructed, and had the government been even mildly adequate, it might have remained a disaster. As things worked out, it turned into a regional catastrophe with perhaps 316,000 dead ( without a census, it’s hard to know the real figur). So we‟d better have a pretty good basic knowledge of earthquakes and the processes that produce them. In this chapter we‟ll study those basic processes, how earthquakes are located and their intensity measured, the terminology used, and how earthquakes are predicted (which is ‘poorly’). For those who’ve already had this background, this will be a quick review chapter. 2.0 Seismic Energy Source We normally accept the definition of an earthquake to be the ground shaking that accompanies sudden movement on a fault; true, it may also be the shaking produced by motion of magma underground, a fast-moving landslide, or even an underground nuclear bomb explosion – but those are very rare. If an earthquake produces ground motion, then energy must be consumed in the process. So, where does the energy come from? Figure 1 shows pretty convincingly that earthquakes (dots show locationare associated with plate boundaries (we considered that before in Chapter 3 and we‟ll consider it again in this chapter). Following our plate tectonics discussions, we can accept that the boundaries of plates are active regions, and most of that action takes the form of:  normal faults,  reverse and thrust faults, and  strike-slip or transform faults (Fig.2). Just because there is a break or fracture through a rock doesn‟t mean that the blocks on either side move readily (although if they do, we call the process ‘creep’). Normally, there is enough friction (or resistance to motion) from the rocks on either side of the break to allow a great deal of stress energy to build ( Press your open hands tightly together and try and move one over the other; you will have to use a fair bit of energy to make the movement – and when you do, the hands move with a „jump). In nature, when the resistance across the break or fracture is overcome, the blocks „jump‟ along the fault plane ( as we demonstrated in Fig.2, a fault is just a fracture across which motion has) to new positions. It is the release of that energy that produces an earthquake. Figure 3 is the traditional sketch expression of the relationship between the stress energy imposed upon a rock and the strain (or deformation) of it. The diagram – the result of experiments in labs – tells us that rocks respond in two different manners when subjected to stress – prior to fracturing, that is. At low applications of stress, the rock will deform but as soon as you release the stress, the deformation disappears; that response is termed „elastic‟. If, however, you impose still greater stress, the rock will deform to the point that, when the stress is released, not all the deformation disappears; that‟s termed „plastic deformation‟. Eventually, obviously, even greater stress results in rock fracture – and, as we said above, that‟s what produces an earthquake because the stress is released suddenly. Bear in mind that cold and very brittle rocks (i.e. those very close to Earth‟s surface) tend to have virtually zero elasticity and break with the application of rather little stress. Rocks that are more deeply buried beneath surface ( and thus warm or even ho) will normally survive more stress before rupturing. A good illustration of the release of energy across a major fault is seen in Figure 4 ( illustration of principle of elastic r); here, the progressive deformation is obvious from the increasing curvature of the fence, and following rupture both sides of the fence again are straight but offset. 3.0 Seismic Energy Expression Let‟s create a hypothetical earthquake – in a nice rural region where there‟ll be no dramatic damage – by release of energy that produces a reverse fault (and the surface evidence is a scarp or cliff on the uplifted side) (Fig.5). It‟s pretty rare (impossible?) to find rock that‟s uniformly homogeneous, so some spot along the fracture plane is weaker than any other spots – and that‟s where the failure (or rupture) will start. That exact point is called the focus (sometimes called the ‘hypocenter; the point at surface exactly vertically above the focus is called the epicenter. The failure on the fault plane will be greatest right at the focus, and will gradually fade out in all directions along the plane as the energy of the release is consumed by moving the rock blocks. The greater the release of energy, the further the blocks move, the larger the fault, and the larger the earthquake. The seismic energy waves travel outward from the focus in all directions, but there are different kinds of motion to those waves, some travel through anything, some get totally bogged down in liquid, and others travel only along surface. The very first indication you‟d have of an earthquake is likely to be the arrival of the „jolt‟ and sound (rumble) of an energy wave called „primary‟ (or P-wave). The P-waves travel as a series of fast compressional waves; as the „front‟ of each wave passes, there is a If you scroll fast, you get a pseudo- animation! momentary compression of atoms or molecules in everything it passes through (Fig.6a). During this time, the more brittle material will probably be cracked. In some regions of California, where people experience moderate earthquakes fairly commonly, there is a well-recognized roar like thunder as the front of a P-wave passes and rock breakage occurs. P-waves travel fast, but are dependent on the density of the material they pass through; in really dense rock, they may travel at 8 or more km/s whereas in less dense rock, the velocity is about 5 or 6 km/s. But no matter the velocity, P-waves will travel though anything. The second event you‟d experience would make it hard for you to stand up – arrival of a series of waves that oscillate up and down (Fig.6b). These are termed „shear waves‟ or „secondary waves‟ (or S-waves) and are rather slower than P-waves (averaging something in the range from 4.5 to 3.5 km/s, depending upon the density of media). You‟d expect them to wreck havoc upon the materials broken by the earlier P-waves – and they do! S-waves will travel though all solids but, because you can‟t shear a liquid, they don‟t get through liquids. [Together, P- and S-waves are commonly called body waves because they are strong enough to be recorded through the body of Earth (except that S-waves can’t get through the liquid core)]. Finally, along come the really destructive events: surface waves. Surface waves are slow (because instead of going through Earth they have to take the long way along the surface), averaging perhaps 2 to 3 (or slightly morkm/s. There are two types of surface waves: Love waves move forward across the surface like a slithering sidewinder snake (side to side oscillation) and Rayleigh waves move forward but in an up and down backward rotating motion (Fig.7) – (which is just the opposite of wave motion on a lake). Because the amplitudes of both types are very large, they are very good indeed at knocking everything in their path (like houses) absolutely flat. 4.0 Measurement: Location and Magnitude These days, we have instruments for every measurement on Earth (and beyond). If we need an instrument to measure earthquakes, we need one that‟s not itself getting „all shook up‟ by the quake, right? But we also need to measure the shaking, so it really has to be on Earth‟s surface. If the surface is shaking after an earthquake, just how do we go about this? Well, what we do is make use of the principle of inertia – resistance to motion. Recognizing that the base of the instrument is going to vibrate because it‟s resting on Earth‟s surface, we attach great heavy masses to the actual recorder part of the instrument, knowing that the property of inertia will give us a short breathing spell before the vibration of the earthquake can get those big masses to move. We call the instruments „seismographs‟ and, although combined into one unit, the two aspects of it (one to measure horizontal motion and the other to measure vertical motion) appear as in Figures 8a and 8b. The seismic signals are recorded on a constantly revolving drum graph, and one strip of such a graph is shown in Figure 9. 4.1 Location To locate the epicenter, you need to record an earthquake on a minimum of three seismographs – a process called triangulation. The simplest method is to measure the time interval between the arrival of the first P- and the first S-waves at each seismograph. Figure 10 shows a composite graph of records from seismographs in Denver, Colorado; St. John, New Brunswick; and Lima, Peru. What is NOT shown here is a bit of a fiddle: the bottom axis of the graph says Distance from Epicenter. But that assumes we know exactly how fast the seismic energy waves move through all the hugely different rocks between the site of the seismograph and the epicenter of the earthquake – which is nonsense! What we do is use a table of “Average Seismic Wave Travel Distance per Unit Time” compiled of experimental results from thousands of observations of the velocity of seismic waves through rock of different composition and density. So the distance is not going to be exact – but hopefully not too wild. Now, using a really good map with a really accurate scale, use a drafting compass, place the needle leg at the point of the seismograph location and scribe a circle of the radius obtained from the distance/time table (noted above) (Fig.11). Repeat that for each of the three seismograph locations and the unique point of intersection is the location of the earthquake epicenter. Of course, this is a task that is usually computerized, so locations can now be determined within minutes of an earthquake anywhere in the world. 4.2 Magnitude You have to admit, to completely quantify an event as complex as an earthquake with a single number is bound to raise problems. But this is what the media-fed public seems to want – so we‟ll have a go at it! Previously, we mentioned Charles Richter – and this is the point where his work becomes paramount. Before Charles Richter created an objective magnitude scale, the strength of earthquakes was measure subjectively: the amount of observed damage, the stories recounted by witnesses, extent of damage to structures whose strengths were known, etc. The scale complied from the accumulation of such evidence was called the Modified Mercalli Intensity Scale ( more about it bel). Useful as it was (and still is), the Mercalli Scale tells little about an earthquake in a region essentially devoid of people and buildings. Charles Richter ( remember that associate of Be), af? seismologist at Cal Tech, believed that intensity could be determined by properly interpreting the strength of energy waves reaching seismographs. By 1935 he had the first objective (empirical) scale devised; not exactly surprisingly, it became known as the Richter Magnitude Scale. In fact, simple in principle but plagued by terrible complexity of application, there now are many different Richter Magnitude Scales, all designed to fit particular parameters. The original magnitude scale was designed to work within 100 km of an earthquake – after all, Richter was most concerned with the events near him – and this is termed the „local magnitude‟ scale, and the designation is M . Lo make things easy to determine M from L the readings of a strip graph, Richter set up a scheme called a nomograph (Fig.12) that consists of three strips of paper such that a straight line joined 100 km (on the distance scale, in log units), 3 (on the arithmetic magnitude scale), and 1 mm (on the log amplitude scale) – and he called that the “standard earthquake”. From there on, it seemed to be a „snap‟ to get the magnitude of any earthquake; all you needed was the distance away from your seismograph (you could get that from the triangulation method), the amplitude from your instrument‟s graph (for whatever set of energy waves you chose to work with), and then read the magnitude from the straight line joining those two points. To fully appreciate that this is a logarithmic magnitude scale, look at Fig.13; a magnitude event of 6 is 10 times more energetic than one of 5! But does the 10 times factor hold all the way through the scale? Certainly from scale numbers up to about 7 it does: each magnitude unit increase means about 10 times more seismic shaking ( i.e. a magnitude 7 event produces 10 times more shaking than a magnitude 6 event and 100 times more shaking than a 5 event); but from magnitude 7 up, the relationship sort of falls apart. The frequency of shaking seems to peak somewhere between magnitudes 7 and 8. From there up, higher magnitude numbers mean the shaking has affected bigger and bigger areas, and the duration of that shaking is much greater too. The end result of course is that with the damaging shaking lasting longer, destruction becomes greater and greater with higher magnitudes. So, the total amount of energy released (and the devastation produced) by a megathrust earthquake of magnitude 9 most definitely is much greater than one of 8. Nevertheless, Richter Scale magnitudes above 8 are simply not reliable ( we say the scale ‘saturates’ at about 8 because any scale reading beyond 8 is invariably smaller than the reali). Just before we look at how this problem of getting accurate magnitude numbers for the largest earthquakes was solved, let‟s just see how well the subjective Mercalli Scale and the empirical Richter Scale estimates compare. Actually, not badly at all (Fig.14)! In fact, while virtually every university sports a seismograph (as do a huge number of high schools), the United States Geological Survey (USGS) still sends out questionnaires to „qualified people‟ all over the world after every major earthquake so that they‟ll have a Mercalli Scale reading just in case there‟s something controversial about the Richter readings. 4.3 Moment Magnitude Scale To remedy the non-scalar behavior of large earthquake measurements, in 1979 a new scale was introduced: the Moment Magnitude Scale, M . Thiw is the scale that Charles Richter would have created if he‟d known how! The scale is based upon the total energy released by an earthquake, thus is applicable to every possible instance, even the very great earthquakes – and it never „saturates‟ or becomes „non-scalar‟. Moment magnitude is not based upon seismograph instrumental recordings of a quake. Rather, total energy is calculated from  total measured area of the fault that has ruptured  amount of offset along the fault (i.e. how far did the blocks move), and  the strength of the rocks involved. Up to about magnitude 7 or so, there‟s no significant difference between Richter Scale readings and Moment Magnitude numbers; from there up, the differences can be significant. For example, a massive earthquake off the coast of Chile in 1960 was interpreted by Richter Scale seismographs to be 8.5; Moment Scale calculations place it at 9.5 – the largest earthquake ever! In fact, it seems likely that 9.5 is just about the maximum earthquake magnitude that can be generated by any plate tectonic process on Earth. We don‟t know of any rocks strong enough to store enough stress energy to give anything higher without breaking. Certainly, as far as scientific research is concerned, the Moment Magnitude Scale now supersedes the Richter Scale ( but for a quick estimate – right after an earthquake – everyone relies on the Richt). Scale I‟m sure you see a great advantage to use of the Moment Magnitude Scale: it allows us to study earthquakes from the distant past just by detailed mapping around fault zones, and measuring offsets. This study of ancient earthquakes by „reading‟ the rock record is called paleoseismology. 5.0 Destruction 5.1 Acceleration, Period and Resonance If you look in a dictionary, the definition of acceleration is: The rate of change of velocity with respect to time. A moving body can accelerate by changing speed or by changing direction. Americans - with their aversion to anything metric - measure acceleration in units of g (i.e. acceleration 2gainst the force of gravity), where 1 g = 32 ft/sec/sec (in metric terms that‟s 980 cm/sec ). Obviously, acceleration has two components: horizontal and vertical. Acceleration in the horizontal sense is very obvious to us; at about 0.1 – 0.2 g of horizontal acceleration we have trouble standing. That‟s about what you‟d experience as a Via Rail train swayed around a sharp bend in the tracks when it was up to speed. At 0.1 g of horizontal acceleration structural damage begins to most buildings. At something in the range of 1.8 g, destruction is pretty much total. Vertical acceleration is not so bad – things go up, things go down! As long as there is no accompanying horizontal motion, structures can suffer quite a high vertical acceleration without coming apart. Here‟s a great example of how factors can suddenly work against you! Let‟s say you are standing beside a tall, wood flag pole that is topped by a bloody great big brass eagle (guess what country you are in!). Reach over and give the pole a push; depending upon its construction, it will move back and forth with a certain period; let‟s say that period is 2 seconds. Now, instead of you shaking the pole, let‟s say there is a seismic wave passing in the ground. No matter what pushes the pole, it will vibrate with a period of 2 seconds (because that’s what the construction/dimensions of the pole dictate). Now, let‟s suppose the seismic wave also has a period of 2 seconds! At this point the pole will resonate in much the same way as if you were pushing your kid on a swing at just those proper moments that will make the swing go higher [This property of resonance explains why certain wavelengths of sound – such as might emanate from a (truly very accomplished) opera singer - can build to the point of shattering a glass]. Well, my whole story builds to the point that we can suggest that the flag pole is sitting in a chunk of concrete, and (because of the dimensions/construction properties of that base) it starts resonating at a 2 second period – and the whole pole just vibrates to pieces! Believe it or not, unless construction engineers pay attention to the properties of the materials they use, that can happen to buildings! Fig.15 is a schematic to illustrate how different materials and poor construction allowances in adjacent buildings in Mexico caused the tops of the buildings to smash against each other in a severe quake in 1985 as seismic waves produced different resonance periods in the different structures. Live and learn! 5.2 Liquefaction Liquefaction can be a fatal process. It‟s a pretty easy process to demonstrate on a small scale; walk at a reasonable pace along the waterline of a sandy beach and you can be confident that the wet sand will support your weight. However, start jumping up and down in the same spot and in seconds you‟ll be up to your ankles in a slurry of sand and water! The sand sediment weight (and some of yours) is partially supported by the water in the pore spaces between the sand grains. When the sand is shaken, the grains are pushed apart, grain-to-grain contact is lost, and the sand- water mixture suddenly behaves as a liquid (such a mixture is sometimes called quicksand). During earthquakes, intense shaking can cause water- saturated sediment to change rapidly from a solid to a liquid. Liquefaction of poorly compacted sediment has caused multi-storey apartment buildings to tip over (Fig.16), highway bridges to collapse and dams to fail. One amazing indicator of liquefaction are „sand blows‟, small mounds of sand formed at the surface where water from liquefied layers squirts out of the ground (Figs.17,18: a pocket-knife serves for scale). 6.0 Plate Boundary Quakes Starting in Chapter 3 and again at the beginning of this chapter, I‟ve emphasized the connection of earthquakes with plate boundaries (take another look at Fig. 1). At this point, let‟s summarize the characteristics of earthquakes to be expected at each type of boundary, according to the stress acting at that boundary. 6.1 Divergent Boundaries At divergent plate boundaries the stress action is tending to pull the opposite sides apart (Fig.19). Think about a slab of cold, hard rock: it won‟t come as a huge surprise that under tensional stress rock is weak. Earthquakes at these locations, therefore, should be numerous, close to surface, and low magnitude – and that‟s what we see in Fig.19. So, rocks breaking apart at a place like the Mid-Atlantic Ridge probably won‟t initiate tremors that worry anyone. Obviously, oceanic crust is quite thin so it breaks with little stress applied (and earthquakes will be low magnitude); continental crust is much thicker, so it takes a bit more stress to pull it apart (and earthquakes will be higher magnitude). The type of fault developed by such stress will likely be normal faults (look at Fig.2 again), where there is an extension to the crust as one side of the slab undergoing rupture slips down a sloping surface. 6.2 Convergent Boundaries The largest earthquakes in the world originate in subduction zones (Fig.20). The largest earthquake ever recorded by seismographs occurred in the subduction trench off the west coast of Chile in 1960 (which we mentioned above and will treat in some detail in a case study, next c), and the second largest appears to be the one from the trench off the west coast of Sumatra on 26 December 2004 (again, look for details in the n). Nearly all the energy released by earthquakes world- wide comes from subduction zone ruptures; as an example, the huge Chile 1960 quake was alone responsible for 30-45% of all
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