Pelagic clay, also known as red clay and brown clay, is a fine sediment found in some parts of the abyssal plain. In this article we shall discuss its origin, its distribution, and its lithification.
Origin and distribution of pelagic clay
The bulk of pelagic clay has its origin as very fine sediment (which is therefore clay in the sense of having small grain size) wind-blown from the land, chiefly fine particles of quartz and of clay in the mineralogical sense; volcanic ash also contributes, and even cosmic dust, i.e. micrometorites fallen from outer space.
We should not be surprised to find sediment from the land being deposited in the middle of the ocean, since the fine dust raised by duststorms can be carried right across the ocean from continent to continent; dust storms in the Sahara, for example, can trigger smog warnings in Florida, and are also the main source of mineral nutrients for the Amazon Basin.
This dust is spread liberally across the surface of the ocean, but as you can see from the map below, not all of the ocean bed is covered with pelagic clay (represented by the brown areas on the map).
This is not because the clay is not deposited in other regions, but because in other areas other sediments are also deposited to the extent that the resulting sediment is considered to be nearshore, siliceous, calcareous, terrigenous, glacial, or turbidite in origin. Pelagic clay is therefore the "none-of-the-above" sediment; it is what you get where there are insufficient nutrients for coccolithophores and foraminiferans, or the sea floor is below the CCD; and where there are insufficient nutrients for diatoms and radiolarians, or the sea floor is below the OCD; and which is sufficiently far from the ablation of glaciers into the sea that glacial sediment is not deposited; and ... et cetera, et cetera.
Pelagic clay is the most slowly deposited of all sediments, typically accumulating at less than 10 mm per thousand years; hence it is easily swamped by other sources of sediment if these are present.
The color of pelagic clay varies from place to place, depending on the source of the dust that composes it. Rich browns or reds are common, these colors being produced by iron-based minerals.
In texture it is of course fine-grained, since only the finest and lightest materials can be wind-borne so far out to sea.
As it is compacted by burial, the flakes of which the clay is composed reorient themselves to lie flat and parallel to one another.
Compared with other muds, pelagic clay is remarkable for having very little in the way of sedimentary structure. The photograph below is of ten centimeters of a core sample of pelagic clay from the South Pacific, to demonstrate how extremely boring its structure is.
Because of the extremely gentle mode of deposition, pelagic clay will exhibit no ripple marks, sole marks, cross-bedding, grading by size, or other sedimentary structures; the most we may see are the faintest suggestions of extremely fine laminae.
Lithified pelagic clay: how do we know?
As with other fine-grained sediments, the lithification of pelagic clay is caused chiefly by compaction.
The resulting rock can be recognized as having its origin as pelagic clay by virtue of its mineral composition, its extremely fine-grained and dense texture, and by its lack of sedimentary structure.
Furthermore, pelagic claystone will be rich in ichthyoliths: disjointed parts of fish, usually teeth and scales, testifying to its marine origin. It may also contain the occasional siliceous or calcareous test; or in some cases it will be interbedded between layers of limestone or chert, serving as another indication of its origin.
In previous articles we have mentioned the sedimentation rates for various sediments. Perhaps it is time to ask ...
How do we know?
There are a number of ways we can find out the present, or at least recent, rate of sedimentary deposition.
One of the simplest involves what is known as a sediment trap, which is used to measure deposition of marine sediment. One can be seen being recovered from the ocean in the photograph below.
Fundamentally, it is a technologically sophisticated bucket. One aspect of its sophistication is that it has a whole set of collection bottles at its base which successively rotate into the collection position at fixed intervals. This allows geologists to measure seasonal variations in the quantity and composition of sediments.
This method is particularly useful when dealing with the very fine sediments with a very low rate of deposition such as compose oozes and pelagic clay.
Another method, suitable for when sedimentation rates is higher, is to take a drilling sample in which some layer corresponds to a recent event. For example, the first tests of hydrogen bombs in the 1950s are marked in the sedimentary record by the sudden appearance of cesium-137, an isotope of cesium not found in nature.
Another such time marker consists of the peak in environmental lead that occurs in 1970. Before that point, lead in sediments rose with the use of petroleum; in 1970, the U.S. Congress passed the Clean Air Act, and the lead found in sediments begins to decline.
Other methods involve a rather ingenious use of naturally occurring radioactive isotopes which are constantly being deposited on the sea floor; we shall describe these methods in detail in the article on the U-Th, U-Pa, and Ra-Pb methods of dating.
When we look at sedimentary rocks, and try to figure out their rates of deposition, we may also appeal to more conventional methods of radiometric dating. If we have (for example) a layer of pelagic mudstone, itself undatable, sandwiched between two layers of igneous rock, which is datable by radiometric methods, then by subtracting the date of the lower layer from the date of the upper layer we get a period for the deposition of the mudstone in between, and so, given the thickness of the mudstone, we can figure out its average rate of deposition over that period. If we wish, we can then use our knowledge of how much more compact pelagic mudstone is than the parent sediment to produce a sedimentation rate expressed in millimeters or the original sedimentary material per thousand years.
There is quantitative agreement between the rates of deposition measured in sediments today and the rate of deposition inferred for the corresponding sedimentary rocks; for example the deposition rates calculated for the calcareous material in marine limestone are the same as the deposition rates measured for calcareous ooze; and similar remarks apply to other rocks and their corresponding sediments.
This agreement is confirmation, if any is still needed, that geologists are correct in their diagnosis of the mode of deposition of the parent sediments of these rocks.
In this article we shall discuss the deposition of glacial till in a marine environment. Readers my find it useful to refresh their memory of the relevant terminology by looking back at our previous article on glaciers.
Deposition of glacial marine sediment
The mechanism by which glaciers form and move has already been discussed in our main article on glaciers.
In that article we discussed the sediment deposited by the ablation of glaciers on land. Glacial marine sediment is formed when instead glaciers ablate into the sea: the rafts of ice carry the glacial till out to sea, where they melt, depositing the sediment.
It follows from their mode of transportation that although glacial marine sediments are composed of the same sort of material as glacial sediment on land, this material will be arranged quite differently: on land, we find the coarser material arranged in moraines, while the finer material is carried off by meltwater and spread across outwash plains. At sea, by contrast, the coarser sediment will be deposited at random among the finer sediment.
Glacial marine sedimentary rocks: how do we know?
We can observe present-day glacial marine sediment in the Arctic and Antarctic. These sediments have sufficiently distinctive characteristics that when we find rocks with the same characteristics in the geological record, we are entitled to deduce that they are lithified glacial marine sediments.
The marine fossils in such rocks identify them as being marine. What identifies them as glacial is their unique structure. This consists principally of finely ground rock flour, often lacking layering, scattered throughout with dropstones: boulders and cobbles which have been rafted out to sea on glacial ice and then sank when the ice melted. The photograph below shows lithified glacial marine sediment particularly rich in dropstones. It is clear from their arrangement that the larger stones have been deposited in the finer sediment from time to time as it accumulated, rather than, for example, being a breccia the gaps in which have been filled with a finer sediment, for in that case the cobbles would be in contact with one another.
Note how the rock flour forming the matrix is clearly the same material, finely ground, as the dropstones.
Being glacial material, dropstones are often unrounded or poorly rounded. Furthermore, some of them will betray clear signs of their glacial nature by being polished and striated as a rest of being dragged along at the base of the glacier.
These features allow us to identify glacial marine sedimentary rocks. Our ability to do so helps us to find out about climatic conditions in the past; we shall discuss this further when we address paleoclimatology.
Saline giants are vast deposits of soluble minerals. What do we mean by "vast"? Well, to take one example, the Louann Salt covers 800,000 square kilometers and is four kilometers deep, amounting to some seven quadrillion tonnes of salt. Even a small saline giant, such as the one found in the Michigan Basin, covers an area of 100,000 square kilometers and has a depth of 250-350 meters.
Although precipitation of salts from seawater can be observed today in an ordinary bucket, the formation of saline giants cannot be observed anywhere. This is not really surprising: the geological record shows that the formation of saline giants has only happened at certain times and in certain places, and it is not unexpected that we should happen to be living in one of the times when no saline giants are forming.
While it is not unexpected, it is annoying. Our understanding of sedimentary rocks is in other cases greatly enhanced by the fact that we can see the sediments being deposited in the present day. In the case of saline giants, we lack this information and must do the best we can.
In this article we shall review what is known, and what may plausibly be conjectured, about the formation of saline giants.
Evaporation of seawater
Seawater contains a variety of dissolved ions, such as (in descending order of abundance by mass) Cl-, Na+, SO42-, Mg2+, Ca2+, K+, HCO3- and Br-; these eight ions alone make up more than 99% of the dissolved ions in seawater, and other ions can be neglected for the purposes of this article.
When seawater evaporates, these precipitate out as minerals such as halite (NaCl) and gypsum (CaSO4·2H2O). The evaporation of 1 liter of seawater will produce around 35 grams of evaporites.
The degree of evaporation required for precipitation varies from mineral to mineral: so gypsum will begin to precipitate out when seawater has been reduced to about 30% of its original volume, but it needs to be reduced to 10% of its volume for the precipitation of halite.
The relevant facts are summarized in the table below. The first column specifies the mineral, the second gives its abundance as a percentage of all the minerals precipitated, and the table as a whole has been ordered roughly according to the ease with which the various minerals precipitate out, from those that precipitate out most readily down to the most soluble. Minerals which occur only in the tiniest traces have been omitted.
The figures given here are based on the pioneering work of Usiglio, still thought to be reasonably accurate: more information will be found here.
The upshot of all this is that as sea water evaporates, a small quantity of calcite will be deposited first. As evaporation continues, gypsum will be deposited: as there is much more gypsum than calcite, and as most of the calcite will have been deposited already, this means that the gypsum will swamp the calcite being deposited, and so the resulting rock will be almost exclusively gypsum. Similarly, when the halite starts being deposited, more halite will be deposited than gypsum or calcite, and the result will essentially be halite. A really intensely concentrated brine, reduced to a few percent of its original volume, will precipitate other salts, but the greater abundance of halite will ensure that it predominates.
We should note also that further dehydration of gypsum (CaSO4·2H2O) will remove the water molecules associated with the calcium sulphate, converting it to anhydrite (CaSO4).
It follows that saline giants will, broadly speaking, consist either of halite, gypsum, or anhydrite.
Models for the formation of saline giants
Saline giants are invariably covered over by a blanket of more conventional sediment, otherwise they would long ago have been washed away by rain (if on land), or dissolved in the sea (if underwater). This observation leads us to a question which is initially perplexing: how in the world can they have formed in the first place? The sea dissolves soluble minerals, and is nowhere near the saturation point at which they must start precipitating out. How, then, is it even possible for these saline giants to form within a marine basin?
So, how were the saline giants formed? Your first guess might be that they are the result of a bit of sea becoming closed off from the main body of the sea and simply drying up. However, there is simply too much salt to be accounted for by a single such event. The drying up of a kilometer's depth of seawater would result in the deposition of only 14 meters of salt. What we need is some model in which the basin keeps being filled, either continuously or periodically, with new supplies of salt water.
We shall describe four such models. Note that although they cannot all apply to the same saline giant at the same time, it is perfectly possible for them to apply to different evaporite formations, or, conceivably, to the same evaporite formation at different times. In this sense, it is possible for all the models described to be correct.
Model 1: A barrier with a small gap in it
Our first model is this: suppose we have a sedimentary basin which is connected to the sea only by a very narrow channel. Combine this with an arid climate and little input of fresh water into the basin from rivers and streams or from rain. So long as the rate of evaporation is greater than the input of fresh water, the physical necessity that the surface of the sea inside the basin must always be at the same level as the surface of the sea outside the basin ensures that salt water will always be flowing into the basin; and as water will be continually evaporating from it, leaving the dissolved minerals behind, this will increase the salinity of the water in the basin until it reaches the saturation point and precipitation occurs.
Something of the sort can be seen today: it was first calculated by Edmond Halley (of Halley's Comet fame) that the Mediterranean loses more water to evaporation than the input of fresh water; this, he realized, explained why there is always a current flowing into the Mediterranean through the Straits of Gibraltar.
We may note that the Mediterranean is indeed somewhat more salty than the Atlantic. However, there is no saline giant forming on the floor of the Mediterranean today. The reason is (so it has been calculated) that in order for this proposed mechanism to work, the cross-sectional area of the channel must be many orders of magnitude smaller than the surface area of the water in the sedimentary basin; otherwise the tendency of water to mix will prevent the water in the basin from ever becoming saturated enough for precipitation to take place. The Straits of Gibraltar today are simply not narrow enough for the formation of saline giants.
Now a channel of just the right size would be unstable: we might expect it within a short space of time to get blocked up or, alternatively, broadened out, either of which would spoil the proposed mechanism. The deposition of the saline giants under the Mediterranean took 300,000 years --- a blink of an eye by the usual standards of geologists, but a long time for such a narrow channel to stay just the right width. For this reason, although we must regard this proposed mechanism as possible in theory, we should require very definite evidence to endorse it in any particular case.
Model 2: A barrier overtopped at high tide
Another proposal is of a barrier (a "sill") between the main body of the sea and a sedimentary basin such that water will surmount it only at high tide. Again, this is possible, but as with the mechanism previously described, it needs to be just right, and, to stay just right for hundreds of thousands of years. This is possible but implausible. A further problem with such a model is this: the influx of water at high tide needs to just balance out the water lost by evaporation. For if the water input was less than the water lost, then the water in the basin would be reduced to an intermittent puddle which would not account for basin-wide sedimentation; and if the water input was greater than the water lost, then eventually the basin would fill up until, at high tide, it was at the same level as the main body of the sea, allowing mixing to take place. It is hard to see what effects could keep the situation in equilibrium, so that the basin is always reasonably full but never quite fills up.
Model 3: A barrier overtopped by rises in sea level
A third, similar model again requires a barrier completely blocking off the sedimentary basin, which is periodically overtopped, not as a result of the tide, but as a result of an increase in the global sea level caused by changes in the Earth's climate. Such a model would predict that layers of evaporites should alternate with layers of more conventional marine sediments deposited during periods when the basin is full. It seems that this applies in the case of the Mediterranean deposits (see here for further information) but it is by no means true of all such evaporites.
This requires a less precise set of circumstances than the previous model, in that fluctuations in global sea level caused by climate change might be expected to be greater in magnitude than local fluctuations caused by the tide.
This model might be combined with the first or second models: variations in sea level might alternately allow and prevent the mechanisms described in model one or two: again, we should then expect to see an alternation of evaporites with more conventional sediments.
Model 4: A permeable barrier
A fourth model is as follows: the basin is completely cut off from the main body of the sea, but by porous sediment or sedimentary rock, so that sea water can seep through the barrier. As with our other models, we require that the output of fresh water through evaporation should be greater than the input: however in this case, unlike the "narrow channel model", the water in the basin is free to drop below sea level when it evaporates: this produces a pressure differential between the two sides of the barrier and ensures that water flows in just one direction, from the main body of the sea into the basin. The nearest analog to this model in the modern world would be the deposition of salt in lagoons.
You will notice that this model does not require anything to be just right: neither the height nor the width of the barrier are crucial: so long as the barrier is above sea level, the system described will work.
Such a model is immune to the problem of equilibrium that we raised with respect to the second model. For the lower the water level sinks in the basin by evaporation, the greater the pressure differential on the two sides of the barrier, and the greater the influx of water; and conversely, the higher the level in the basin, the less water will seep through the barrier. So we might well expect such a system to be in equilibrium, with the basin never either drying out completely or filling up so as to overflow the barrier.
Saline giants: what do we know and how do we know it?
All the models described above require two things: that the basin in which the saline giant is deposited should be nearly or totally isolated from the main body of the sea; and that the climate should be such that more fresh water is lost through evaporation than is input by rivers and rain. We can test whether these conditions were in place, and show that these are the conditions under which saline giants form. For example, the Mediterranean is in the present day nearly cut off from the sea, and it does not strain the imagination to suppose that 5.9 million years ago it was more isolated still. To take another example, it would be strange to see evaporites forming in the Gulf of Mexico today; but conditions were just right at the time when they formed, when it was almost, or entirely, blocked off from the main body of the sea by what is now West Africa. To take a third example, the Castile formation in the Delaware Basin (which, despite its name, is located in Texas and New Mexico) was ringed around by a reef complex during the time of its formation. Independent evidence from paleoclimatology also confirms that the climatic conditions were right for evaporite formation in each case.
So although in many cases there is controversy over which of the models we have described best explains the existence of a particular saline giant (a controversy which perhaps in some cases will never be fully resolved) this is really a dispute about details: geologists are agreed, and can confirm, that what is required is a basin which is almost, or entirely, cut off from the main body of the sea, plus an arid climate.
This explains why such deposits are rare in the geological record. There is no particular reason why this set of circumstances should be common: they occur by happenstance and not by any sort of geological inevitability.
In order to fully understand plate tectonics and the evidence for it, it is necessary for the reader to know a little about the physical properties of rocks. In this article we provide a brief introduction to the concepts involved.
Stress and strain
In ordinary English, stress and strain are more or less synonymous. In physics, they refer to different though related quantities.
Stress is a measure of the force per unit area exerted on a surface of a deformable body. That is, roughly speaking, stress is to solids what pressure is to gasses, and like pressure it is measured in pascals (Pa); that is, in newtons per meter squared.
Strain is the deformation of a body as a result of stress. In geology, strain is given by the length by which a rock expands or contracts divided by the length it was originally: because this is the ratio of a length to a length no units are associated with it.
Tension, compression, and shear
The stress on a rock (or any other material, for that matter) can be classified as tension, compression, or shear, as illustrated in the diagram below.
Rock is strong under compression but relatively weak under tension and shear. This is a result of the microscopic structure of rock: it contains microscopic cracks which are forced open and enlarged by tension and shear but which are forced closed by compression.
This is why a small overhang on a cliff will easily break under its own weight (being subjected to shear) whereas the rock at the foot of the same cliff will bear the much greater weight of all the rock above it, as in that case it is being subjected to compression.
Elastic and plastic behavior
A material is said to be elastic if it recovers from stress --- that is, if, having been bent or extended or compressed under shear stress or tension or compression, it snaps back into its original configuration when the stress is released.
A material is said to be plastic if, on the contrary, once stress has squeezed it into a certain shape, it retains that shape; plasticine, for example, is plastic at room temperature and surface pressure.
When a solid is placed under stress, its behavior is at first elastic; then (with increasing stress) plastic; then with the addition of enough stress it fractures.
A solid which undergoes very little plastic deformation between elastic behavior and fracturing is said to be brittle. In colloquial English we usually reserve this work for things which are both brittle in the technical sense and also require little stress to break, such as egg-shells; in its technical use in physics, however, a substance such as diamond is also brittle in the technical sense: diamond may not break easily, but it will break before it undergoes any significant plastic deformation.
The opposite of brittle is ductile.
A material will have greater resistance to fracture if it is under a high surrounding pressure; and it will be more ductile at higher temperatures.
The reader should also bear in mind that the rate at which stress is applied may be significant: a force rapidly applied may produce fracture which, if more slowly applied, may produce deformation. The material known as Silly Putty is famous for clearly demonstrating this property: it deforms under gentle pressure from one's fingers but shatters if hit with a hammer.
Application to rocks
We should now explain how all this applies to rocks in particular.
Rocks on the surface will exhibit elastic and brittle behavior, since they are cold and at low pressure. At depth, the pressure will be greater, increasing their brittle strength (that is, their resistance to fracture) and the temperatures will be higher, decreasing their ductile strength (that is, their resistance to plastic deformation).
Below the depth at which the ductile strength is less than the brittle strength, the rocks will be fully ductile and plastic. Some people describe the rocks below this depth as molten, but this is not accurate: they are not a liquid, but rather a ductile solid, like plasticine.
As a result, rocks near the surface tend to fracture under stress creating geological faults; more deeply buried rocks tend to fold.
Earthquakes are also a phenomenon of the upper, brittle, elastic part of the rock. When two pieces of the Earth's crust try to move past one another, their mutual friction impedes them and they bend very slightly. Earthquakes are caused when the potential energy of the bent rocks is sufficient to overcome the resisting friction and they snap back, releasing the stored energy in the form of kinetic energy. This is only possible if the behavior of the rocks is elastic rather than plastic. Consequently we do not expect (and do not find) deep earthquakes except when they are associated with subduction (as will be discussed in a subsequent article).
How do we know?
The behavior of rocks at surface temperatures and pressures are easy to verify. To find out how they would behave at the greater temperatures and pressures requires special equipment.
An example of the results of one such experiment can be seen in the photograph below. The samples in this experiment were three cylinders of marble, originally identical and looking like cylinder (a).
* Cylinder (a) has not been subjected to any stress; it is a control sample.
* Cylinder (b) has been subjected to 20% strain with 27 MPa of confining pressure (where 1 MPa = 1,000,000 Pa). This is the sort of pressure you'd get at a depth of about 1 km. As can be seen from the photograph, it has undergone both plastic deformation and brittle fracture.
* Cylinder (c) has been subjected to 20% strain with 46 MPa of confining pressure: the sort of pressure you'd get at a depth of about 1.75 km. As you can see, it has deformed without any fracture as though it was made of putty.
More recent experiments have reached greater levels of sophistication. By using lasers to heat rock samples, and a device known as a diamond anvil cell to exert pressure on them, it is possible to simulate temperatures and pressures such as are found deep within the Earth. The photograph below shows the setup for a diamond anvil cell experiment at the Argonne National Laboratory.
Such methods do not tell us everything we would like to know. Reproducing conditions in the very core of the Earth would require some sort of breakthrough in material technology. Another thing that is hard to simulate is the effect of time. We know that materials are more likely to deform and less likely to shatter when stress is applied gradually: so what happens if you apply a gentle stress to a rock over a period of millions of years?
Such questions can to some extent be answered with reference to established notions in physics; but clearly if all such questions could be answered with total accuracy with reference to purely theoretical considerations, then geologists wouldn't spend so much money on diamond anvil cells and lasers.
That being said, what we do know is sufficient for us to understand plate tectonics; certainly it is quite enough for an introductory course such as this one.
In this article we shall review some important facts about the physics of seismic waves (that is, waves generated by earthquakes).
This article is background reading for the following article on the structure of the Earth, which is itself background reading for subsequent articles on plate tectonics.
Surface waves and body waves
When an earthquake occurs, it is the cause of seismic waves, including both waves that travel along the surface (Love waves and Raleigh waves) and waves that travel through the body of the Earth (P-waves and S-waves, known collectively as body waves).
In this and succeeding articles we shall be interested only in body waves, since by traveling through the Earth they give us clues about the Earth's interior.
P-waves and S-waves
P-waves are waves of compression and tension, like sound waves; indeed, they are sound waves, traveling through rock rather than through air. S-waves are waves of shear: that is, of displacement at right angles to the direction of travel of the wave, resembling the waves produced by shaking the end of a rope.
The image above shows the motion of P-waves (top) and S-waves (bottom). Look carefully at the picture of P-waves. If you focus on any particular vertical line, you will see that it is merely oscillating from side to side: it is the regions of compression that are moving from left to right, while the medium itself has no net motion. Similarly in S-waves, it is the displacement that moves from left to right, while the medium through which the waves move exhibit no such motion.
If we use κ to represent the incompressibility of a medium, μ to represent the rigidity, and ρ to represent the density, then the velocity of P-waves through that medium will be given by
and the velocity of S-waves by
From these formulae we can immediately see that a P-wave will always travel faster than an S-wave through the same medium. We can also see that S-waves will not travel through liquid at all, since liquid has no rigidity and so in the case of liquids µ = 0.
Refraction and least time
The principle of least time says that a wave traveling through a medium will take, not the shortest route as measured by distance, but rather the quickest route between two points.
The word refraction means the change of direction undergone by a wave when it passes from a material which permits travel at one speed to a material which permits travel at another speed. The existence of refraction is a direct consequence of the principle of least time: in the generalized diagram of refraction shown to the below, the path of the ray is the quickest route from point A to point B.
Using the principle of least time, we can say exactly how refraction should take place: a little simple mathematics tells us that the ratio of the velocity in medium A to the velocity in medium B should be equal to the ratio of sin θA to sin θB; or to put it another way, equal to the ratio of xA/dA to xB/dB. This is known as Snell's Law. Where the velocity changes smoothly and gradually through the medium, this will result in the ray taking a curved path.
The upshot of all this is that if we know the wave velocities associated with each point in an object then we know exactly how a wave will travel through it, since its motion is determined by the principle of least time.
When a wave encounters a sudden transition between mediums with different associated velocities, so that the wave is refracted at an angle, some of the energy will also be reflected back in a direction determined by the well-known law that the angle of incidence is equal to the angle of reflection (see diagram, below). This is why one can see faint reflections in windows and in water.
P-waves and S-waves behave in the same way as more familiar waves such as light, but with one difference: light is exclusively a transverse wave, and reflects back as one. When either a P-wave or and S-wave is reflected, however, the reflection will be composed of both P and S-waves.
When such reflections of P and S-waves are detected, this tells us that they are being reflected off some sort of sharp boundary between rocks having different physical properties and hence different associated velocities.
How do we know?
The properties of body waves can be studied in the laboratory. They can also be derived from more basic principles, simply working from the fact that they are waves and must therefore obey the physics of waves. (Note that in physics a "wave" is not just something that behaves in a kind of wave-y sort of way, but something the dynamics of which can be described by the wave equation. The mere fact that body waves are waves therefore tells us quite a lot about them.)
As discussed, the routes taken by P and S-waves through the interior of the Earth are determined completely by the velocities at which these waves travel at each point in the interior.
If these velocities were the same for every point in the Earth, then the time it took for a wave to travel from the focus of an earthquake (the focus being the point where the earthquake occurs, not to be confused with the epicenter, which is the point on the surface directly above the focus) to an earthquake detector (a seismometer) would be proportional to the distance of a straight line drawn from the one to the other through the body of the Earth.
But this is not the case. By studying the data showing how long the waves do take to travel through the Earth, it is possible to determine the velocities of P-waves and of S-waves at each point in the Earth.
Doing so is what mathematicians know as an "inverse problem": it might be compared to trying to reconstruct the shape of an object by observing its shadow. It would obviously be much easier to work the other way round and deduce the shadow from the object; and similarly the general problem of seismic tomography --- that is, of discovering the inner structure of an object by studying the passage of body waves through it --- would be very difficult to solve.
But fortunately we do not have to solve this problem in the general case, but only for one particular object: the Earth. This problem has a property that makes it particularly easy to analyze.
Consider the fact that the time it takes for a P-wave or an S-wave to get from the focus to a seismometer depends to a good degree of approximation only on the angle of separation between the focus and the seismometer. This tells us that (again to a good degree of approximation) the values of vP and vS at any particular point within the Earth must depend only on the depth below the surface of the Earth and not on the longitude and latitude (in technical terms, we may say that the values of vP and vS are spherically symmetric).
Consequently, what looked like a three-dimensional problem involving finding values of vP and vS for every depth, longitude, and latitude can be reduced to the one-dimensional problem of finding values of vP and vS at every depth.
This fact makes it possible to produce a graph such as that shown below, relating vP and vS to temperature; the figures are taken from the Preliminary Reference Earth Model (here). One fact that you should notice immediately is that the speed of vS drops to 0 in the outer core, showing that it is liquid.
We have said that the travel time of body waves depends on the angular separation between the earthquake focus and the seismometer to a good degree of approximation. By studying the small variations from this rule, it is possible to detect low and high velocity anomalies within the Earth: volumes where the waves travel slower or faster than what we should expect if the Earth's interior was perfectly spherically symmetric. Work on this is ongoing.
I hope the physicists round here will tell me if I've got anything wrong, and the non-physicists will tell me if I've been unclear.
Well, I'm just a library administrator but I do have a degree in geological engineering with an emphasis on geophysics. Unless the laws of physics have changed since I last payed attention, you are dead on correct in both facts and explanation so far as I know.
Edited by anglagard, : change misplaced semicolon to apostrophe
Read not to contradict and confute, not to believe and take for granted, not to find talk and discourse, but to weigh and consider. - Francis Bacon
I note that there are velocities for the S waves in the inner core. It seems to me that the outer core would filter out the S waves - The S waves could not get to the inner core.
So far as I can see, they S-waves can't get into the core and they can't get out. But they can still exist in the core because when P-waves partially reflect off the boundary between the inner and outer core, they partially partially reflect as S-waves, as mentioned in the article.
Even if there were no S-waves ever in the core, it would still be meaningful to say what vS was in the core, because it can be derived from two properties which definitely are in the core, namely rigidity and density. It would be how fast S-waves would travel though the core if there ever were any.