3.3. This includes the types of decay processes, their half‐lives, and the useful age range for dating by each of these methods.
Of the three decay series, the most commonly used is238U → 206Pb. This is because238U is much more abundant than the other two radioactive isotopes and thus it and its daughter product are easier to measure accurately. Typically, heavy minerals, such as zircon, sphene, and/or monazite, are used for analyses because they contain substantial actinides, are relatively easy to separate and resist chemical alteration.
A problem arises because some206Pb occurs naturally in Earth materials, without radioactive decay, so that measured238U/206Pb ratios include both daughter and nondaughter206Pb. This must be corrected if an accurate age is to be determined. The correction can be accomplished because yet another lead isotope,204Pb, is not produced by radioactive decay. The ratio of204Pb/206Pb in meteorites, formed at the time Earth formed, is known and is taken to be the original ratio in the rock whose age is being determined. The correction is then accomplished by measuring the amount of204Pb in the sample and subtracting the meteoritic proportion of206Pb from total206Pb to arrive at the amount of206Pb that is the daughter of238U.
In most analyses, the ratio of235U/207Pb is also determined. The235U/207Pb decay series has a much shorter half‐life than the238U/206Pb series, so that the U/Pb ratios are much smaller for a given sample age. Once the ratios of both parent uranium isotopes to the corrected daughter isotopes (238U/206Pb,235U/207Pb) have been determined, sample ages are generally determined on a concordia plot (Figure 3.14) that displays the expected relationship between the reciprocal ratios206Pb/238U and207Pb/235U as a function of sample age.
Figure 3.14 Uranium–lead concordia plot (blue) showing sample ages as a function206Pb/238U and207Pb/235U ratios and samples with lower206Pb/238U and207Pb/235U ratios (red), reset during a 2.5 Ga metamorphic event.
Either decay series is capable of producing a sample age on its own. For example, a206Pb/238U ratio of 0.5 indicates an age of approximately 2.6 Ga (Figure 3.14), while a207Pb/235U ratio of 40 indicates an age of approximately 3.7 Ga. One advantage of these decay series is that the measurement of multiple, closely related isotopes permits ages to be checked against each other and provides robust sample ages when both methods yield a closely similar age. For example, if a sample possesses a206Pb/238U ratio of 0.7 and a207Pb/235U ratio of 32, a sample age near 3.5 Ga (Figure 3.14) is highly likely. When two age determinations are in agreement, their ages are said to be concordant. Another way to state this is that samples that plot on the concordia line when their206Pb/238U versus207Pb/235U ratios are plotted yield concordant ages. Such dating techniques are especially powerful in determining robust crystallization ages of igneous rocks that crystallized from magmas and lavas (Chapters 8 and 9).
A significant problem arises from the fact that subsequent events can alter the geochemistry of rocks, so that age determinations are no longer concordant. For example, when rocks undergo metamorphism, lead is commonly mobilized and lost from the minerals or rocks in question. This results in lower than expected amounts of daughter lead isotopes, lower Pb/U ratios and points that fall below the concordant age line on206Pb/238U versus207Pb/235U diagrams. When samples of similar real ages, affected differently during a single metamorphic event, are plotted on a concordia diagram, they fall along a straight line below the concordant age line. One interpretation of such data is that the right end of the dashed straight line intersects the concordant age line at the original age of the sample (4.0 Ga in Figure 3.14) and that the left end intersects the concordant age line at the time of metamorphism (2.5 Ga in Figure 3.14). Wherever possible, such interpretations should be tested against other data so that a robust conclusion can be drawn. One such independent radiometric decay series, with widespread application, is described in the section that follows.
Rubidium–strontium systematics
Several isotopes of the relatively rare element rubidium (Rb) exist; some 27% of these are radioactive87Rb. Most Rb+1 is concentrated in continental crust, especially in substitution for potassium ion (K+1) in potassium‐bearing minerals such as potassium feldspar, muscovite, biotite, sodic plagioclase, and amphibole. Radioactive87Rb is slowly (half‐life = 48.8 Ga) transformed by beta decay into strontium‐87 (87Sr). Unfortunately, the much smaller strontium ion (Sr+2) does not easily substitute for potassium ion (K+1) and therefore tends to migrate into minerals in the rock that contain calcium ion (Ca+2) for which strontium easily substitutes. This makes using87Rb/86Sr for age dating a much less accurate method than the U/Pb methods explained previously.
One basic concept behind rubidium–strontium dating is that the original rock has some initial amount of87Rb, some initial ratio of87Sr/86Sr and some initial ratio of87Rb/86Sr. These ratios evolve through time in a predictable manner. Over time, the amount of87Rb decreases and the amount of87Sr increases by radioactive decay so that the87Sr/86Sr ratio increases. At the same time, the87Rb/86Sr ratio decreases by an amount proportional to sample age. A second basic concept is that the initial amount of87Rb varies from mineral to mineral, being highest in potassium‐rich minerals. As a result, the rate at which the87Sr/86Sr ratio increases depends on the individual mineral. For example, in a potassium‐rich (rubidium‐rich) mineral, the87Sr/86Sr ratio will increase rapidly, whereas for a potassium‐poor mineral it will increase slowly. For a mineral with no87Rb substituting for K, the87Sr/86Sr ratio will not change; it will remain the initial87Sr/86Sr ratio. The87Rb/86Sr ratio of the whole rock however decreases at a constant rate that depends on the decay constant.
In a typical analysis, the amounts of87Rb,87Sr, and86Sr in the whole rock and in individual minerals are determined by mass spectrometry, and the87Rb/86Sr and87Sr/86Sr ratios are calculated for each. These are plotted on an87Sr/86Sr versus87Rb/86Sr diagram (Figure 3.15). At the time of formation, assuming no fractionation of strontium isotopes, the87Sr/86Sr in each mineral and in the whole rock was a constant initial value, while the87Rb/86Sr values varied from relatively high for rubidium‐rich minerals such as biotite and potassium feldspar to zero for minerals with no rubidium. These initial87Sr/86Sr and87Rb/86Sr values are shown by the horizontal line in Figure 3.15. As the rock ages,87Rb progressively decays to87Sr, which causes the87Sr/86Sr
