The Uranium Fission Clock

Drawings of mass-spectrometer charts showing the isotopic spectra of two kinds of strontium: (A) common strontium and (B) strontium from an old mineral rich in rubidium.(Seepage 32for photo of a mass spectrometer.)

Rubidium-87 and strontium-87 fall on the same spot in the mass spectrum. Therefore, rubidium must be separated chemically from strontium before the strontium can be analyzed in a mass spectrometer. It is done with ion-exchange columns. Four of them are shown in this photograph. The author of this booklet is adding a sample, dissolved in a few drops of hydrochloric acid, to the second column.

Let us take as our closed system a mica crystal in a mass of granite. Mica contains a fair amount of rubidium, and it retains its radiogenic strontium very well. Furthermore, mica crystals are often associated or even intergrown with the slender, rod-shaped crystals of a mineral called apatite—a phosphate of calcium. It is justifiable, on the basis of geological knowledge, to say that the mica and the apatite grew at roughly the same time and thus presumably from the same liquid medium that became granite when it later solidified. Now strontium is geochemically similar to calcium, and some strontium will have gone into the apatite crystal in place of calcium. Apatite contains no alkalis—hence apatite will have virtually no rubidium (which is an alkali) in it to contaminate the ⁸⁷Sr. Consequently, when we find apatite in an old granite, we know the apatite will still contain the kind of common strontium that was taken into the mica crystal when it grew originally.

We can separate the apatite from the granite by standard mineralogical techniques, extract the strontium from theapatite chemically, and analyze it on a mass spectrometer to obtain the isotopic spectrum—the relative amount of each isotope that is present. We can then perform the same isotopic analysis on the strontium extracted from the mica, and subtract the original (apatite) strontium from the total (mica) strontium, to obtain the radiogenic component or daughter product. (Seepage 32for details of this method.)

Obviously, there is a certain error associated with every isotopic analysis, so such a calculation is meaningful only when the radiogenic component is large compared with the error in the measurement of isotopic abundance. When one large quantity must be subtracted from another large quantity to obtain a small difference, there is an obvious limit to how much one can trust the result. The absolute accuracy in measuring strontium isotope abundance is a few tenths of 1%, using the best mass spectrometers now available. In practice, one can trust a calculated age for a specimen only when the ⁸⁷Sr is as little as about 5% radiogenic. The results do not mean much when only 1 or 2% is radiogenic.

A sample of granite being made ready for crushing and mineral separation.

When a neutron strikes the nucleus of uranium-235 (²³⁵U) or plutonium-239 (²³⁹Pu), it may cause the nucleus to split into two roughly equal fragments, releasing neutrons and energy. This is the well-known process of neutron-induced fission, the method in which nuclear energy is produced in both reactors and bombs.[14]The most common uranium isotope, ²³⁸U, also breaks up by fission, but does so all by itself, without the need for any external neutrons. That process isspontaneous fissionand it goes on at random, very much like radioactive decay. It is a relatively rare process and the fission half-life is long—about 10 million aeons (10¹⁶ years). That means that only about one spontaneous fission occurs in uranium-238 for every 2 million alpha decays. That is enough to make a useful clock, however, because ²³⁸U is present almost everywhere. (See Table III onpage 19.)

Imagine an atom of ²³⁸U in some mineral. When the atom suddenly fissions, it breaks in two with considerable energy, and the two fission fragments rip like cannon balls through the surrounding crystalline structure in opposite directions, creating havoc along the way. They travel a distance something like 10 microns (4 millionths of an inch) before they are finally slowed down and stopped by all their collisions with other atoms. Each fragment’s path remains behind as an intensely damaged tube through the crystal.

The process was known for a long time before anyone was able to find these fission tracks (the damaged tubes) in the crystals. Finally, about 1960, three young physicists, R. L. Fleischer, P. B. Price, and R. M. Walker, working at the General Electric Research Laboratory, fell upon the idea of etching freshly broken surfaces of crystals with acid. They reasoned that a region so intensely disturbed by the passage of a fission fragment should be etched more easily and deeply than the undisturbed surrounding crystal. That idea turned out to be correct, and fission tracks have now been found in almost every common mineral (since almost all minerals contain small amounts of uranium).

Tracks of uranium fission from a fossil antelope bone fragment from Hopefield, Cape Province, South Africa.

The fission clock method works this way: A cleavage face or a polished surface of a crystal or glass fragment is etched with a suitable solvent. Different acids work best for different materials, and a suitable procedure must be developed especially for each substance. The etching brings out the fission tracks so they can be seen (usually as little conical pits) and counted under a microscope.

After this, the sample is exposed to a known amount of slow neutrons[15]in a nuclear reactor. New fissions are produced, but this time only in ²³⁵U (which is present in all natural uranium in the proportion of 1 atom of ²³⁵U to 137.7 atoms of ²³⁸U), because slow neutrons do not produce fissions in ²³⁸U. After the neutron irradiation, the same surface is etched again, and the new tracks counted. The old tracks, having been etched twice, now appear larger and thus can be distinguished from the new ones that were caused by ²³⁵U fission.

The rate at which ²³⁸U decays by fission, λf, is known, as are the rate it decays by alpha decay, λα, and the total number of slow neutrons,n, to which the sample was exposed in the reactor. The age of the crystal or glass can then be calculated:

t =1λαln (1 +nNsNi× constant)

and the constant has the value:

½λα× 582 × 10⁻²⁴λf× 137.7= 4.25 × 10⁻¹⁸

Fission-track dating is a brand new technique, still only partly developed. It has enormous range and is applicable to numerous minerals; these advantages imply that it is likely to become very useful.

An atomic absorption spectrophotometer is used to measure the amount of potassium in samples of mica dissolved in acid.

The most complicated and therefore probably the most interesting decay scheme of all is the decay of uranium to lead, discovered well over half a century ago and still intensively studied. There are several reasons for the interest.

First, uranium and lead are geochemically separated to a high degree, not only on the small scale of an ore deposit but also on the scale of the earth as a whole. Second, natural uranium has two isotopes with half-lives that are neither too long nor too short to be useful (the greater half-life almost exactly equaling the age of the earth), and these half-lives differ from each other by a factor of about 6.3. That leads to very important consequences, as we shall see. Third, uranium and lead are both common, and techniques are available for extracting them in measurable quantities from almost any natural material.

As a consequence of these happy circumstances, the study of uranium and lead has contributed a great deal to understanding the earth’s history and the processes that go on inside it. F. G. Houtermans, one of the great pioneers in this study, jokingly called the methodPLUMBOLOGY, and it seems a useful name.

The greatest achievement of the “plumbologists” has been the calculation of the age of the earth, first proposed by Houtermans, a German physicist, and independently by Arthur Holmes, a British geologist, in 1946 and finally perfected by C. C. Patterson in 1953. It is actually a rather simple calculation, although the way to discovering it was far from easy. Before we look at it in detail, however, let’s consider some basic assumptions and explain what is meant by “the age of the earth”.

From studying the mechanics of the solar system, scientists have become reasonably certain that the earth and the other planets and their satellites all were formed in a common process in a relatively short period of time, geologically speaking. Perhaps it took a dozen millionyears or so, but compared to the time that has elapsed since, that is a twinkling. At some time soon afterwards, the earth became molten, or at any rate fluid enough to allow much of its iron to settle toward the center to form the earth’s core. Similar cores presumably formed in other planets. As the iron went down, it took some lead with it, and as the silica went up, uranium followed it toward the surface, because of the chemical affinity between these kinds of elements. In the present earth, we have found, almost all the uranium is concentrated in the top layer, or crust, which is only about 25 miles thick under the continents and even thinner under the oceans.

Internal structure of the earth. The central core is probably an alloy of iron and nickel, surrounded by a mantle of less dense silicate material, with a thin crust of still lighter silicates.

The time of this early and relatively rapid separation of uranium and lead on a worldwide scale is the event that plumbologists can determine, and the period since then is what they mean by “the age of the earth”. When Houtermans first wrote about it, he called it “the age of uranium”.

How is this done? We have said that one of the isotopes of uranium, ²³⁵U, decays faster—about 6.3 times faster—than the other, ²³⁸U. They decay into two different isotopes of lead. Therefore, if we can determine the isotopic composition of average ordinary lead in the earth’s crust today, and if we can somehow obtain a sample of the kind of lead that is locked in the earth’s core, we can calculatehow long it took to change thePRIMORDIALlead (like that in the core) into present-day lead in the crust by the gradual addition of radiogenic lead—lead that has resulted from the decay of uranium. Now, someone might logically ask, “Isn’t it necessary to know also the actual amount of uranium involved in the process, and isn’t this difficult to determine?” It turns out to be a remarkable aspect of the Holmes-Houtermans calculation that the uranium-concentration terms cancel out in the equations and only theratioof the isotopes and their decay constants need be considered. These are all known accurately.

Next, we must decide just what is average present-day lead? It isn’t enough to go to a lead mine and get a sample, because, unfortunately, leads from different mines have widely varied isotopic composition—that is, a different mixture of four natural isotopes, ²⁰⁴Pb, ²⁰⁶Pb, ²⁰⁷Pb, and ²⁰⁸Pb—as a result of their geologic histories. No, lead samples from a mine won’t do. However, geologists have been able to separate lead from recent marine sediments, obtained from the ocean bottom, far from land. These are of uniform composition, and are good samples of what the world’s rivers bring into the ocean. Other useful samples can be found in plateau basalts, which are enormous bodies of dark volcanic rock that make up the bedrock in many parts of the world. The lead from these basalts is isotopically very much like the lead in the oceans.

Very well, but how about the lead from the core? Where can we hope to find a sample of it? It turns out to be easier than you might think. Astronomers believe it highly probable that most meteorites are fragments of a former planet that broke up for reasons that are not entirely clear. It is pretty definite, however, that this protoplanet (or these protoplanets, for there may have been more than one) had an iron core, and this core (or these cores) is the source of the iron meteorites sailing around in space. A large meteorite hit the earth not too long ago (geologically speaking) and caused the Meteor Crater near Canyon Diablo in Arizona.

Examining ocean-bottom sediments obtained by lowering a tube-like instrument that brings up a long rod-shaped “core”, prior to nuclear age determination of the sample.

Many fragments of the meteorite iron have been found around the crater, and it is reasonable to assume that this is the kind of iron we would expect to find in the core of the earth. Like the core iron, it is mixed with a little lead, which can be isolated and analyzed in a mass spectrometer for its isotopic composition. This lead is found to be much less contaminated with radiogenic lead, and hence is much more primitive than the oldest leads found on earth. Thus, meteorites presumably are as close as we can get to true primordial lead—the lead of the time when the earth (and the protoplanet) first formed.

Once these measurements were available, it was easy to write the Houtermans equation for present-day and primordial leads in this way:

(²⁰⁶Pb²⁰⁴Pb) present - (²⁰⁶Pb²⁰⁴Pb) primordial(²⁰⁷Pb²⁰⁴Pb) present - (²⁰⁷Pb²⁰⁴Pb) primordial

= 137.7(eλ238t- 1)(eλ235t- 1)

The present ratio of ²³⁸U to ²³⁵U is 137.7.

Substituting the best experimental lead isotope ratios into the equation and solving for t, Patterson was able to calculate that the earth is 4550 million years (4.55 aeons) old. Subsequent calculations based on other procedures generally have confirmed that result.

Each method of nuclear age determination involves a different sequence of sample preparation. Wood, peat, charcoal, bones, or shells are cleaned for carbon-14 dating in order to remove every trace of possible contamination by modern carbon as well as extraneous old carbon. Rocks are crushed and ground, minerals are separated according to what is needed in any particular study, and the desired elements are extracted and separated by chemical procedures. Often there may be several different ways of doing the same thing; different laboratories use different procedures. In every case, however, long and complicated procedures must be followed before results are obtained from which an age can be calculated. There is no such thing as a black box into which you can throw a rock and read its age on a dial!

Of all the elements that are part of the useful parent-daughter systems, only potassium is common enough to be analyzed by conventional chemical techniques. All the other elements, especially the radiogenic ones, are present in such small quantities that special processes had to be developed to measure them. The most valuable and generally used process is calledISOTOPE DILUTION.

This is a process for analyzing an unknown material by incorporating uniformly into it a small amount of a radioactive test substance and determining how much the tracer radioactivity is altered by dilution in the original material.

It works like this: Let’s say that we have an unknown number of atoms,x, of a given element. The normal isotopic composition of this element is accurately known, as it is for most elements, and the ratio of two of its isotopes can be expressed as A/B. We now add toxa known (but usually smaller) amount,c, of the same element. This quantity has a drastically different isotopic ratio, A′/B′. We mixxandcthoroughly together. The ratio A′/B′ can have almost any value, but must be different from A/B and we must know exactly what it is. (There are many ways of determining this chemically, or we can use a sample isotope of known composition obtained from the U. S. Atomic Energy Commission’s Oak Ridge National Laboratory at Oak Ridge, Tennessee.) The substance added is known colloquially as theSPIKE.

After the original material and the spike are thoroughly mixed we have:

x(A/B) +c(A′/B′) = (x+c) (A″B″)

in which A″/B″ will be the ratio of the two isotopes in the mixture. With this information in hand, we can perform any chemical purification or transfer process with the material (see photo onpage 22), without having to worry about loss. (Even if 90% of the material should be lost in some operation, the isotopic composition would not be changed, and that is all we are interested in.) Now we can place the material containing the isotopic mixture in aMASS SPECTROMETER, which will determine the ratio A″/B″. When we have that, we may substitute the value of A″/B″ in the equation and quickly calculatex, the unknown concentration of atoms in the original sample.

A large (12-inch) mass spectrometer (at left) in use. Electronic equipment (right) charts results (seepage 21).

Essential parts of a mass spectrometer. Atoms to be analyzed are changed to ions in the source. Then the ions are accelerated by high voltage, deflected in a magnetic field according to their mass, and the intensity of the separated beams is measured in the collector.

The mass spectrometer measures isotopic abundances using a magnetic field to sort electrically charged particles into groups according to their masses. It works this way: A small drop of material to be analyzed is placed on a metal filament and dried. The filament, in its holder, is placed inside the mass spectrometer, and heated electrically in a vacuum, like the filament in a light bulb. As the wire begins to glow, some of the sample begins to radiate, or “boil off”, losing an electron or two in the process. In other words, some of the atoms will be changed into positiveIONS.

An alternative method is to introduce the sample material into the vacuum chamber in the form of a gas (like argon, for example), and then bombard the gas with electronsstreaming from a hot filament. The electron stream will knock some electrons off the gas molecules and this also will produce positive ions. Either process of ion production is satisfactory, depending on the problem to be tackled, but the mass spectrometers for the two methods are naturally quite different.

Whichever way the ions were produced, they are next exposed to a strong electric field, accelerated, and electrostatically focused into a beam. These charged particles are directed into a magnetic field between the pole faces of an electromagnet. The magnet does the analyzing by the principle of magnetic deflection that was known to André Ampere and Michael Faraday more than a century ago. Any moving electric charge has a magnetic field associated with it. This field interacts with the field of the analyzing magnet to impress a deflecting force on the charge. The force acts at right angles to the direction the charge travels and also at right angles to the direction of the impressed magnetic field. The pull of this force depends only on the electric charge and the speed of each particle: A light single-charged particle will be deflected more than a heavier particle with the same charge. In this way, the ions in the beam are sorted out into a number of separate beams, each made up of particles of the same charge/mass ratio. Each beam contains one isotope of the original material, because isotopes differ on the basis of their mass. By adjusting the current in the electromagnet we can direct these separate beams into a “collector” and electrically measure their intensity one by one. This gives the relative abundance of the separate isotopes in the sample.

Measuring age by one of the long-lived radioisotopes requires a closed system. Usually this is some kind of crystal formed in a period of time that is short, compared to the time that has elapsed since, and that has remained unchanged since it formed. Specifically, neither the parent isotopes can have been added nor the daughter isotopes removed by any process other than radioactive decay.

The earth is a dynamic system, however. Things are always changing and moving—not very rapidly, perhaps, but fast enough, in geologic time, to raise mountains and shift oceans. Solutions are moving around, dissolving something here and depositing it again somewhere else. Temperatures are changing as one place is denuded by erosion and another area buried under layers of sediment. Under such conditions, few systems remain closed. It is perhaps surprising that we find any closed systems at all. Let us look at a few that are known to be reliable. (They are listed in Table I onpage 4.)

In the early 1950s, when the potassium-argon (parent-daughter) method was being developed by scientists at the University of Chicago, it was thought that the potash-bearing variety of the mineral feldspar would be an ideal closed system, because it was usually optically clear and free of flaws. This widely shared, logical, and perfectly scientific deduction soon turned out to be quite wrong. The scientific workers discovered that when feldspar and mica from the same rock (and thus of the same age) were analyzed side by side, the mica always came out older! Investigation showed that feldspar “leaked” argon (lost some of its radiogenic argon) even at room temperature, but the mica retained all or nearly all of the argon that had been generated in it.

With the development of the rubidium-strontium (parent-daughter) method by L. T. Aldrich and his co-workers at the Carnegie Institution of Washington, came the realization that mica was also very useful for this analysis, for it usually contains ample rubidium and not much original strontium that would mask the presence of the radiogenic strontium. As a result, mica, especially black mica (the mineral biotite), has enjoyed great popularity as a good and easy-to-find closed system.

A scientist making adjustments on an “argon train”, a maze of glass tubing in which argon is released from minerals and purified for analysis.

Everything has its limits, and mica is no exception: Even mica tends to leak argon at elevated, but still relatively low (geologically speaking), temperatures. These effects also depend on pressure and other factors, not all of which are well known; these elevated temperatures, pressures and other conditions of course act to some extent on all rocks buried in the earth’s crust. It is known that at only about 300°C at moderate pressures argon is leaked from mica faster than it is being generated in it by the decay of radioactive potassium. The temperature needed to cause the rapid loss of strontium from mica is not much higher. Mica, especially biotite, will recrystallize and lose all its radiogenic constituents (argon and strontium) at temperatures where many other minerals show little or no change.

That means that we cannot always rely on mica to give the date of theoriginalcrystallization of a rock—the time when it cooled from a molten state. Instead, mica will tell us when the rocklastcooled from, say, several hundred degrees centigrade, regardless of what may have happened to the rock before that. The mica may have been reheated as a result of being buried under a few miles of sediment, for example. The mica will show when the rock last cooled—in other words, when it came up again.

In spite of early disappointments with potash feldspar for argon dating, some of it is useful for rubidium-strontium procedures. It all depends on how much original strontium the potash feldspar contains. Most feldspars, unfortunately, contain far too much, but rapid screening by X-ray fluorescence or flame photometry methods can weed these out and identify specimens low enough in original strontium to be useful. Otherwise, feldspar is an excellent closed system for rubidium and strontium; it remains closed even at temperatures high enough to melt many other minerals. It is not affected at all by the same degree of heating that will drive argon out of biotite. The rubidium-strontium age of feldspar usually comes close to the time of original crystallization of the rock.

Obviously, here is a geologically important tool. If we find feldspar and biotite in one rock, and if feldspar, tested by the rubidium-strontium method gives the same age as biotite tested by potassium-argon decay, then we can say with confidence that the rock has not been reheated since shortly after it crystallized. Conversely, if the biotite comes out much younger than the feldspar, we can be sure that somethinghashappened to this rock long after it first crystallized. Such information is not only valuable to pure science—it can also be useful in locating areas favorable for ore prospecting and in other practical ways.

Another very interesting mineral is zircon (a silicate of zirconium), one of the accessory minerals found in small quantities in many crystalline rocks. Zircon usually occurs in very small grains and is heavy and hard, so that it can be separated from the other rock without much difficulty, even though it may take 100 pounds of rock to supply a gram of zircon.

Zircon usually contains a fair amount of uranium and very little lead. It holds radiogenically produced lead well, even at relatively high temperatures. But that is not all. Even if some of the lead is lost, there is a mathematical way of correcting for it. This technique is calledCONCORDIA ANALYSISand was developed by G. W. Wetherill, a physicist then at the Carnegie Institution of Washington. It is based, again, on the fact that natural uranium has two long-lived isotopes—²³⁸U and ²³⁵U—and that the lighter one, ²³⁵U, decays faster than the heavier. The daughter products of both uranium decay processes are isotopes of the same element, lead—²⁰⁸Pb and ²⁰⁷Pb, respectively. Heavy isotopes are not separated to any significant degree by chemical processes, so that if radiogenic lead has been lost from a system for any reason, the other lead isotopes also will have been lost in whatever proportion they were present originally.

If we plot a graph of the radiogenic ²⁰⁶Pb/²³⁸U ratio against the radiogenic ²⁰⁷Pb/²³⁵U ratio for concordant (closed) systemsof all ages, we obtain the curved line shown in the figure below. The curve is the locus ofallconcordant U-Pb ages and is calledConcordia. Then if we test two or more particular zircons of the same age that have lost different amounts of lead, at about the same time, the plot of their ²⁰⁶Pb/²³⁸U ratios against their ²⁰⁷Pb/²³⁵U ratios will fall on a straight line that is a chord of the Concordia curve. The upper intersection of this chord with the curve then will mark the true age of the zircons. This is an elaborate technique utilizing difficult chemical procedures, but it has proved invaluable in solving some important geologic problems.

The Concordia curve offers a useful way of analyzing results of age determinations on the mineral zircon.

The mineral hornblende provides another useful system. Hornblende is a complex silicate of sodium, calcium, iron, magnesium, and aluminum, and usually contains a few tenths of 1% of potassium. It is unusual in that it tenaciously retains its radiogenic argon, even at relatively high temperatures.

Still another good system is the rare feldspar, sanidine, which is excellent for both potassium-argon and rubidium-strontium age determination. Sanidine usually is found in volcanic ash falls and has been important in the establishment of the geologic time scale, as we shall see.

Finally there is still another way of obtaining a closed system by using the whole rock, not just a crystal of a single mineral within it. A large body of granite or similar rock may contain a number of minerals, some or none of which may be closed systems. Yet as long as this body of rock remains impermeable to solutions (which in nature means mostly to water), no substance will be able to move very far in it because diffusion in solids is so slow. Consequently it will remain a closed system, as a whole, regardless of what happens to the individual mineral grains.

If we take a piece from near the middle of this body of rock and if this piece is much larger than the largest constituent grain in it, then we have a fair sample of a closed system—the whole rock. The only difficulty arises from the fact that few rocks are sufficiently impermeable to solutions to retain argon, and many rocks contain so much common strontium that rubidium-strontium analysisis impractical. Still, we can use the rapid survey methods as for feldspar, selecting the few rocks that would be useful. This work has been done frequently, and the results have been fruitful for rubidium-strontium analysis. The whole-rock rubidium-strontium age dates the time when the rock became impermeable.

Fossil skull of Zinjanthropus, nearly 2,000,000 years old, discovered in 1959 by Dr. L. S. B. Leakey in Olduvai Gorge. Accurate dating of this earliest human ancestor was possible by using the potassium-argon method.

One of the most talked-about age measurements in recent years was the determination of the unexpectedly great age of fossil ancestors of man, found by the Britishanthropologist, Dr. L. S. B. Leakey, in Olduvai Gorge in Tanzania. The measurements were made by Garniss H. Curtis and Jack F. Evernden at the University of California in Berkeley by the potassium-argon method. The age came out a little less than 2 million years, about twice as old as it “should be” in the view of many scientists. Human remains of such great antiquity had never been found before, and much doubt was raised about the validity of the figures.

Time periods as short as two million years are not easy to measure by potassium-argon. The amount of argon produced in that time is extremely small, and contamination by argon from the air is a serious problem. Still, the measurements were repeated, the rocks were studied again, and the result did not change: The fossils were still about 2 million years old.

In cases like this, one tries to find some other method to check the results in an independent way. After many attempts it was discovered that the same rock strata dated by potassium-argon also contained some pumice—a porous volcanic glass—and that this glass was suitable for uranium fission-track dating. The measurements were made in the General Electric Research Laboratory. What was the result? Just about 2 million years!

When such altogether different techniques give the same number, one can have some confidence that the number is exact. It would be difficult to imagine a disturbance in nature that would cause these unrelated methods to give the same wrong number—in both cases by a factor of two. The double check simply means the Olduvai man is 2 million years old. There is not much doubt about it.

Much of historical geology is based on a relationship called theLAW OF SUPERPOSITION. This simply means that when some rock formation was placed on top of some other formation by natural processes (sedimentation or volcanic eruption, for example), the layer on top must be younger than the one on the bottom. Such a conclusion may now seem obvious, but the concept was not even expressed until the very end of the eighteenth century and was still amatter of scientific controversy when Abraham Lincoln was a boy. It was the law of superposition, however, that led the early geologists to establish the first geologic time scales and to realize the enormous extent of geologic time.

Diagram illustrating the law of superposition. Each rock bed is younger than the strata under it.

In essence the system of establishing age by this concept is this: Somewhere a large and easily recognized layer of sedimentary rock was known. It had a characteristic color, texture, gross composition, and overall appearance. Let us call itbed M. This bed could be traced across the countryside until a place was reached where it could be seen thatbed Mrested on another, different layer of rock, which we might callbed L.Bed Lcould also be traced some distance and ultimately could be observed resting on a still different stratum, which we shall callbed K. Some of these beds had fossils in them, and it was eventually realized that rocks with the same kinds of fossils are of the same age, even though they may differ in other respects—in color or composition, for instance.

Ifbed Mwas followed in another direction, perhaps a point was reached where it dipped down a little, and here, was found still another layer—call itbed N—on top ofM. Obviously, the sequence of beds from oldest to youngest wasK-L-M-N. Their relative ages were now established. Over the years, hundreds of geologists described various rock layers and identified the fossils in them. By the middle of the nineteenth century, this “layer-cake” structure of sedimentary rocks was well-known in western Europe. (One may say, parenthetically, that America was geologically a vast unknown at that time. Today we know much more about the geology of Antarctica than anyone a hundred years ago knew about the geology of the United States.)

A system of nomenclature for the “layer-cake” was developed and refined. Gradually this nomenclature was accepted internationally. Long-range correlations between beds of the same age, distant halfway around the globe from each other, were made possible as the science ofPALEONTOLOGYdeveloped. The relative age of almost any rock containing even poorly preserved fossils could be determined anywhere in the world with precision. That is, the age of rock layers in relation to one another was known. But the real age—the absolute age—remained unknown until knowledge of radioactivity provided the necessary clocks.

A geologist carefully maps the layers of soil in which ancient flint tools have been found. Exact correlation of the tool-bearing strata with material datable by carbon-14 analysis must be accomplished before the age of the tools can be determined.

Even this process wasn’t entirely without problems. The difficulty lay in the fact that ordinary sedimentary rocks (shale, sandstone, and limestone) cannot be dated by the usual nuclear methods because they present no suitable closed systems. Only some volcanic sediments can bereliably dated by the mica, feldspar, and zircon they contain; but these ancient ashfalls are rare, usually are only a few inches thick, and are not easy to identify at the surface because they weather quickly to clay. Only about a dozen volcanic beds have been accurately dated in North America. The time points they established are mainstays of the present geologic time scale, but they have had to be supplemented by indirect information.

Steeply dipping sandstone strata from the Cretaceous Period near Gallup, New Mexico. This photo, taken in 1901 (note the horse-drawn wagon), was made on an early government survey of Western lands.

Fossils of bird tracks in sandstone in Death Valley, California.

The most important of these indirect time points are furnished by what geologists callBRACKETED INTRUSIVES. It often happens in geologic history that a mass of rock becomes molten at a great depth and forces its way up through several layers of sedimentary rocks. The sedimentary layers are usually bent and twisted (folded) by the upthrust, and where the cooler rock comes in contact with the molten mass, cooler material is burned (recrystallized). Geologists call that processCONTACT METAMORPHISMbecause the sedimentary rock forms are changed, or metamorphosed, to have another form or composition. Contact-metamorphosed rocks, in spite of the damage they have suffered, may contain recognizable and accurately datable fossils. Thereby the metamorphic rock establishes a lower limit for the age of the intrusion: It must be younger than the fossils in the youngest of the metamorphic rocks it touched.

Now on top of the intrusive rock, we may find another sedimentary rock, deposited on top of the intrusive after it cooled and was exposed by the erosion of overlying materials, perhaps millions of years later. This new sediment may also contain fossils and thus furnish an upper limit for the age of the intrusion. The measured age of the intrusive rock thus can be used to set upper and lower limits on the absolute ages of the two sediments. If we are lucky, the two sediments will bracket a relatively short interval, making our measurement quite precise.

Idealized sketch of a bracketed intrusive. The igneous (molten) rock must be younger than the sedimentary rock (A) it intrudes, and older than the rock (B) that overlies it. The relative age of the sedimentary beds is known from their fossils.

Facsimile reprint of the famous time scale proposed by Arthur Holmes in 1959.

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