How Waves Reveal Internal Structure of the Earth.

Transcription

1 How Waves Reveal Internal Structure of the Earth. Unless otherwise noted the artwork and photographs in this slide show are original and by Burt Carter. Permission is granted to use them for non-commercial, non-profit educational purposes provided that credit is given for their origin. Permission is not granted for any commercial or for-profit use, including use at for-profit educational facilities. Other copyrighted material is used under the fair use clause of the copyright law of the United States.

2 We ll need to remember three of the wave behaviors we have learned about: 1) refraction 2) reflection, and 3) s-waves inability to move through liquids. The fourth behavior (difference in rate of propagation) will come up in another slideshow.

3 Remember that we mentioned earlier that waves reflect off of density discontinuities. The various layers you ve heard about the crust, mantle, and core are made of different materials and have different densities. Their boundaries, in other words, are simply density discontinuities. And so it is possible to see these boundaries as reflections on a seismogram. Unfortunately, as the example at right shows, seismograms are very complicated. There are numerous reflections, refractions, aftershocks, and so on sending waves toward the seismograph. Consequently, reflections from a discontinuity are typically much easier to spot after you already know to expect them. Most discontinuities are first discovered in another way by Curious ways the waves refract. seismograph B source A C Crust (density ~2.5) Mantle (density ~3.5) R I

4 Finding the Moho The Crust/Mantle boundary

5 The boundary between Earth s crust and mantle is called the Moho or M-discontinuity or Mohorovicic discontinuity after its discoverer. There is, on any seismogram, a jiggle that records the first reflected wave from the reflection off the Moho of seismic waves. Somewhere in the noise is another first arrival the first arrival of a wave that refracts into the mantle. The diagram shows how this wave moves (green arrows). When the wave arrives at the discontinuity and hits that boundary it sends waves outward in every direction from where it hit. One of these, represented by the heavy green arrow in the mantle, will travel along the discontinuity. As it goes, some of its energy is returned across the Moho at every step, radiating upward. (Do you see some hint here of why the seismogram is so noisy?) seismograph source Crust (density ~2.5) Mantle (density ~3.5) X X

6 At some distance (X) on either side of the mid-way point between the source and seismograph the wave entering the mantle will send out a refracted wave toward the seismograph that will be the first arrival refracted wave. Notice that this wave travels farther than the reflected wave, and so you d expect it to take longer. But once in the mantle it also travels faster than the reflected wave, so the lag between them is not what you d calculate based on distance alone. For stations close to the epicenter the reflected wave arrives first and there is some lag between it and the refracted wave. But you should be able to predict that the lag time will decrease the farther from the source the station lies. At some distance (called the critical distance) both waves arrive simultaneously, and beyond that the refracted wave actually outruns the reflected wave and arrives first. seismograph source Crust (density ~2.5) Mantle (density ~3.5) X X

7 Mohorovicic, of course, did not know this up front. He had to reason his way through it after noticing a strange pattern in arrival times on seismographs. For many earthquakes and many seismographs he plotted out the distance from the epicenter to the station. The location of the station was known, of course, because they are exactly where people put them. Locating the epicenters is something we ll look at separately, in lab. Suffice to say that he knew where both things were for many, many earthquakes. The relationship between distance traveled and time traveled is generally very straightforward. It s usually among the first things you learn in a physics course, and the graph illustrates it. Assuming a constant speed distance traveled = speed/time. Assume it s about 40 miles to Albany and you can drive it at an average (or a constant) rate of 40mph. Your travel time is 1 hour (40 miles = 40 mph/1h). If you can go 60 you ll make it faster in 2/3 of an hour (40 minutes) (40miles=[60mph/(2/3 hour)]. The graph is linear. The longer you drive at a given speed, the farther you ll go. T I M E DISTANCE TRAVELED BY CAR

8 Correcting for differences in the rocks the waves went through, Mohorovicic could make a good estimate of the average velocity of the waves through the crust, and he predicted that he could also, therefore, calculate the arrival time at any seismograph. The same graph also applies to this it works for cars on highways, waves in the crust, anything that moves. T I M E DISTANCE TRAVELED BY WAVES What he saw was different from what he predicted. His hypothesis about predicting arrival times was wrong, but wrong in a peculiar and interesting way. This was a new observation upon which to base and test a new hypothesis.

9 Dr. M s predictions worked fine for stations within a certain distance of the seismograph a distance he called the Critical Distance. This is shown as segment A on the graph. However, beyond the critical distance, instead of arriving at the predicted times (segment B ) the first waves to arrive came earlier than expected (segment C ). (Look at the graph and make sure you get the implication. The lower slope indicates that the first arrival wave has travelled a greater distance (horizontal axis) in a shorter time (vertical axis). That is, it arrived faster than expected. (Imagine Dr. M scratching his head How could a wave speed up?). Any ideas? T T I I M E E A B Critical Distance C DISTANCE TRAVELED BY WAVES X

10 Finally (or maybe quickly, I don t know) he realized that a wave cannot speed up without moving into a denser medium. He realized that the faster wave, arriving first beyond the critical distance, must not be the same wave as the slower one arriving first within the critical distance! It must be an entirely different wave that had moved through denser material. His hypothesis has been tested and supported many times. For one thing, once the existence of the denser material was discovered it became an easy matter to pick out its reflections on seismograms, which are always there. Incidentally, the wave that was predicted to arrive first does, in fact, eventually arrive, but it is lost in the noise of the seismogram. C B? T T I I M E E A B Critical Distance C DISTANCE TRAVELED BY WAVES X

11 Let s think it through by comparing the original picture to the graph. The red arrow to the right is always part of the path followed by A on the graph. If the station is closer than the critical distance then the red arrow on the left is a continuation of A, but it will not be the first arrival wave beyond the critical distance. In other words it is the B part of the graph beyond the critical distance. T I T M I E M E A (B) Critical Distance DISTANCE TRAVELED BY WAVES X seismograph source (B) A Crust (density ~2.5) Mantle (density ~3.5) X C X

12 C is the complete path followed by all the green arrows. Beyond the critical distance B arrives sometime after C has arrived and we see C first. That wave has followed a longer path, but has been able to move enough faster through the mantle to outrun A and B. B C So the existence of the change in density between the crust and mantle was discovered because of the existence of the critical distance. T I I M E A Critical Distance We can also learn something interesting from how far that critical distance is. What? DISTANCE TRAVELED BY WAVES X seismograph source (B) A Crust (density ~2.5) Mantle (density ~3.5) X C X

13 Remember that we are talking about observations of arrival time of wave. What controls when, say, an automobile leaving Americus will get to some destination? That will depend, of course, on 1) how far away the destination is, and 2) on how fast the car is driven distance and speed. It s the same with waves. Distance and speed dictate how far they travel. If the distance is known the only variable is speed, and vice versa. Alternately the two can be entered into a set of equations and both can be determined (2 equations with 2 unknowns in this case). We do not know up front what the depth to the Moho is, though we can interpret it pretty easily once we know to look for a reflected wave. The exact distances that the waves travel depend upon this and we must either solve for it or determine it from wave reflections. The depth to the mantle is somewhat variable, but we ll come back to this. That done, we can solve for something else. We know that the refracted wave must move faster than the reflected one for that part of its course that lies within the mantle. How much faster it can go depends on the density of the mantle. Knowing both the critical distance and the depth to the moho it is an easy matter to solve for the density of the mantle. (Having some idea that it is made of a rock called peridotite before you begin doesn t hurt either.) So Dr. M didn t just find his discontinuity, he also figured out how deep it sits and what the density of the mantle must be. Not surprisingly, it has the same density as peridotite!

14 Characterizing the Crust

15 We ve mentioned that the depth to the Moho (and therefore the thickness of the crust) is variable, but that variation is not without pattern. In general the crust is thinner in the oceans, off the continental shelves. In most places it is about 5km thick, but under volcanically active regions like the ocean ridges, or parts of the Pacific called large igneous provinces (LIPs) it reaches 10km. Under individual highly active volcanoes it can be even thicker. The dashed line in the diagram parallels where the land surface would be if the Earth were a perfect sphere. Its level is not just arbitrarily chosen it is a calculated equilibrium level for the crust. In every place in the oceans (except under very active rapidly building volcanoes) it lies 1/7 of the way between the surface and the Moho. The high ridges and LIPs have a depth to the Moho beneath them in proportion to how high they stand off the surrounding ocean floor. It appears that they float in the mantle at some appropriate level, like blocks of wood of different thicknesses floating in water. 1/7 of the crust is above the equilibrium line and 7/8 beneath it at any point. (Again, except under especially active volcanoes.) As we ll see next, the continents do something similar.

16 The same general pattern holds under the continents but with a couple of differences in detail. Because the continental crust is less dense the equilibrium level is not 1/8 : 7/8 but ¼ : ¾. In other words, the continental crust floats higher than the oceanic crust like an empty (low density) boat floats higher than a laden (high density) one. Incidentally, we can use this difference in the equilibrium levels to compare the differences in density of the continental and oceanic crusts. Secondly, the continental crust is several times thicker than the oceanic crust. It floats higher like two blocks of wood of different density float at different levels. (Think of oak and balsawood.) So the land elevation stands higher than the seafloor because it is less dense (and floats higher) and thicker (and floats higher). That s why the water is in the ocean why we have oceans. Notice that the mantle is much deeper under mountains than lowlands, and even deeper under bigger mountains than small ones. We say they have roots. (Very active mountains like the Himalaya have roots that are not as deep as expected. They are not sinking to their equilibrium level as fast as they are growing.

17 Matching the scales of the two preceding diagrams and joining them we see that the same equilibrium level applies to both. The oceanic crust is both thinner and denser so it sinks lower below sea level and water naturally fills that basin. The continental crust is both thicker and less dense so it floats too high to be covered with seawater. This is why we have land and sea. Sea Level Crust (continental) Equilibrium Level Moho Crust (oceanic) Moho Mantle This tendency for the crust to find some equilibrium level is called isostasy. It explains several interesting places where the crust has recently moved directly upward or downward, or, in fact, is still doing so. WE will come back to the idea a few times over the next two terms. The glaring question, of course, is what does the crust float in? Presumably the Mantle, but the mantle is made of rock, right? Let s move on

18 Characterizing the Mantle

19 Finding the Moho involved studying arrival times at stations fairly close to the epicenters. Stations farther away are of progressively less use, but we do learn interesting things about the mantle from them. One is that waves do not move according to the linear graph we used above, but according to a curve instead. They seem to get faster with distance. This is partly an illusion and partly real. The curve below shows the measured arrival times of the first p-wave at seismographs up to 10,000 km away. Something weird happens past that and we ll see what that is shortly. 20 Travel Time (minutes) Distance of Seismograph from Focus (km)

20 d d The illusory part is pretty simple to grasp. The distance on the graph s x-axis is the surface distance d between the focus and the seismograph. But it isn t surface waves we see on a seismogram km away, its body waves. Our first guess might be that the distance the wave actually travels is the shortest route between the two points the straight-line distance (d in the picture). But it turns out that that distance (or the speed implied by that distance) is not correct either. The waves arrive a bit earlier even than that wavepath suggests. 20 Travel Time (minutes) Distance of Seismograph from Focus (km)

21 D! d d Once again refraction explains the discrepancy. Wavefronts that go somewhat deeper outrun the straight path because they cross into progressively denser material (and refract downward) on the way in and then into less dense material (and refract upward) on the way out of the Earth. One curved path (D!) will balance distance and speed and outrun all the other paths, including the straight-line path, and arrive at the seismograph first. (The others will show up as noise.) This is just a reminder that we ll be seeing curved paths, and why. 20 Travel Time (minutes) Distance of Seismograph from Focus (km)

22 Before we leave the mantle there are a few loose ends to tie up. 1) It s density increases with depth. This is the best and most consistent explanation for the first arrival times from distant earthquakes. The density is about 3.5 g/cm 3 near the Moho and increases to about 5.7g/cm 3 at its lowest part, as indicated by wave arrival times. 2) The density does not, apparently, increase perfectly consistently. There do seem to be some minor discontinuities within the mantle. The significance of these receives a lot of attention and debate these days, but doesn t concern us much at this level. The best guess is that they mark places where certain minerals can (and do) change into denser versions of themselves. 3) The uppermost part of the mantle is very rigid and brittle as brittle as the crust. Below a depth of about 100km (usually) this gives way to a less rigid material that is plastic able to bend or even flow very slowly like silly putty, only a lot slower. This part of the mantle is called the asthenosphere. This is what the crust floats in. The slow rate of movement is why a mountain chain s roots don t keep up with its roof. How deep the plasticity extends is a matter of a lot of debate, largely because it is part of one hypothetical explanation for what causes tectonic plates to move. If the asthenosphere is very thick, the mechanism works, if it is very thin, the mechanism cannot work. This will come up again in Geology II when you talk about plate tectonics.

23 Finding the Core The Core/Mantle boundary (G-discontinuity)

24 The curve we have just seen that describes arrival times through the mantle only works if the seismograph is within about 103 of arc (~11461 km) of the focus. (That s why the graph stops at ~10,000 km). Both p-waves and s-waves show this odd characteristic, but what happens beyond is different in the two wave types. NOTHING DOING Cutaway View of Earth s Insides.

25 Beyond 143 (15910km) from the focus the p-waves return, but not at the expected arrival times. Clearly their velocity has not been what we d predict for the mantle, so they must have moved through material of a different density. (The exact wave behaviors are very complex in the real Earth. I have simplified these drawings to make what s happening apparent without tracing out exactly how the waves behave.) P-waves return Cutaway View of Earth s Insides.

26 P-waves arrive where the green dashed arrows are, but not where the red zones are. (There are some but their timings are very screwy and they are almost unnoticeable on a seismogram.) This region of no p-waves (shown in red) is called the p- wave shadow zone. Make sure you realize that this picture is an idealized picture of Earth cut exactly in half. Cutaway View of Earth s Insides.

27 If we look at the Earth from the outside viewing the surface instead of the interior the p- wave shadow zone would be shaped like a doughnut. The black dot is exactly 180 of arc away from the focus. The blue regions would receive strong p-wave signals, the shaded zone would not.

28 These observations were first explained by a geophysicist named Gutenburg. He realized that the waves at some depth must cross into an even deeper layer than the mantle with a noticeably higher density. Gutenburg could determine the depth to this discontinuity (~2900km) by noticing where the shadow zone began 103 from the focus. That s where the curved wavepaths can just graze it and keep going to where they are expected. The next wavepath over hits it, gets refracted, and winds up somewhere else. Gutenburg realized that where the waves pick up again depends upon how much they have been refracted what the angle of refraction is, in other words. That, remember, depends on the difference in density. Remember that the lower mantle has an estimated density of ~5.7 g/cm 3 based on the bending of the waves moving through it. Given this it is a simple matter to determine the density of the lower layer. Again, the model you see here is very much oversimplified, but you should be able to dig the basic logic out of it. The wave refracts into the inner layer, then refracts again on returning to the mantle and so does not go where it s expected. Cutaway View of Earth s Insides.

29 The deeper layer is called the core, of course. The boundary between it and the mantle is called the Gutenburg discontinuity or simply G-discontinuity after the one who explained its existence and characteristics. Cutaway View of Earth s Insides.

30 Like the p-waves, the s-waves do exactly as expected up to a distance of 103 from the earthquake. Beyond that, no s-waves at all are recorded. This is the s-wave shadow zone.

31 If we look at the Earth from the outside the s-wave shadow zone covers nearly half the planet! The black dot is exactly 180 of arc away from the focus. The blue regions would receive strong s-wave signals, the shaded zone would get none. What can explain this? Certainly not refraction.

32 Characterizing the Core

33 Complete blockage of s-waves by the core indicates that it is liquid (that is, molten), at least in its outermost part. LIQUID! This also explains the odd change in velocity of p-waves, which slow down even though they have moved into a denser material.

34 Remember that the G-discontinuity lies at about 2900km beneath the surface. The angle of refraction of waves across the horizon (entering the core from a mantle of density = ~5.7 g/cm 3 ) suggests that the density of the outer core is about 9.9 g/cm 3. Inge Lehman first suggested that the weak p-waves received within the p-wave shadow zone were consistent with reflections from horizon at a depth of ~5200km, and this was quickly verified and agreed upon. The implication is that the innermost part of the core is solid rather than liquid. ~2900km Outer Core Inner Core Thus we infer an outer core (liquid) and an inner core (solid). Pressure in the core probably means that the density at the base of the outer core is close to 12 g/cm 3. The inner core is presumably solid because of the extreme pressure at that depth. Because of this the inner core density is greater, probably reaching 13 g/cm 3. ~5200km

35 If we sum up all the known and estimated densities of the various layers, and if we take into account the volume of material in each one, then the overall density of Earth can be calculated at around 5.5 g/cm 3 right where Newton put it. Thus we see the test of our last hypothesis the inside of Earth must be made of denser stuff than the outside. Not only do we learn from earthquakes that this is so, we learn what the various internal parts are, what the depths are to them, and how dense each one is. This is quite a lot of information about something we cannot even see!

36 This slide summarizes the important points about the internal structure of Earth. Be able to label the parts and tell their basic characteristics liquid or solid, composition (rock type) and general (relative) density. Don t worry about exact density values. Remember that there re two kinds of crust. If we opened the Earth, took out a pieshaped piece, and enlarged it