Oxygen Isotopes - A GEOL 1122 supplemental reading

Basic information about oxygen isotopes

      Oxygen consists of three stable isotopes: ¹⁶O, ¹⁷O, and ¹⁸O. Most of the world's oxygen is ¹⁶O, a little is ¹⁸O, and a tiny proportion is ¹⁷O. ¹⁷O is so scarce that we can ignore it and focus on ¹⁶O and ¹⁸O.

      This means that a mineral sample with oxygen in it (e.g., a piece of calcite in a fossil) has a measurable ratio of ¹⁸O to ¹⁶O. If we measure that ratio, it's very small, so a mathematical transformation is used to turn that ratio into a more easily managed number. The resulting number is a δ¹⁸O or "delta 18 Oh" value. Values of δ¹⁸O for minerals usually range between -10 and +10. The larger the number, the more ¹⁸O the sample has; the smaller the number, the less ¹⁸O the sample has.

      So who cares? The answer is that the oxygen isotope composition of a mineral sample (its δ¹⁸O) depends on the temperature at which the mineral forms and the oxygen isotope composition of the water from which it formed. Both of those things tell us about ancient environments. Here's why:

1. Temperature dependence of mineral formation

      When an oxygen-bearing mineral precipitates chemically from an aqueous solution (from seawater, for example), oxygen gets segregated. Some oxygen atoms go into the newly-formed solid mineral, and many remain in the solution. The heavier oxygen atoms (the ¹⁸Os) vibrate more slowly and move more sluggishly, and so they are preferentially included in the mineral (where atoms are trapped and can't move as freely).

      That segregation might not mean much to us, except for this: The extent of the preferential inclusion of ¹⁸O in minerals is dependent on temperature. The colder it is, the more ¹⁸O is preferentially included in the mineral that is forming. Thus, if two samples formed from the same water (e.g., seawater), the one with a higher δ¹⁸O formed at a lower temperature. That's a relative statement; in fact, the math exists to calculate an absolute temperature of mineral formation if one knows the mineral's isotopic composition and the isotopic composition of the water from which it precipitated.

2. How glaciation affects oxygen isotopes in the ocean

      When seawater evaporates from the ocean to make rain and snow on land, oxygen isotopes play a role in determining which water molecules evaporate and which don't. A water molecule (an H₂O) in which the oxygen atom is an ¹⁶O is lighter and vibrates faster than a water molecule in which the oxgen atom is an ¹⁸O. Evaporation therefore favors water molecules with ¹⁶O, so that water vapor in the atmosphere is ¹⁶O-enriched. As clouds pass over land and rain falls from them, the heavier water molecules (the ones with ¹⁸O) tend to form the rain, so that the remaining vapor is even more ¹⁶O-enriched or ¹⁸O-depleted. By the time atmospheric water vapor reaches the poles and falls as snow to make glacial ice, it is very ¹⁶O-enriched or ¹⁸O-depleted.

      The result of this is that glacial icecaps store ¹⁶O-enriched or ¹⁸O-depleted water. That means that the ocean's water (the residue of water left after formation of those icecaps) has to be ¹⁶O-depleted or ¹⁸O-enriched during times of extensive glaciation. If we measure the δ¹⁸O of a fossil from the ocean (i.e., a fossil that formed from the ocean's water), that number can tell us the extent to which glaciation had depleted ¹⁶O from seawater or enriched ¹⁸O in seawater - it can tell us the extent of the formation of glacial ice.

3. How glacial ice serves as a record of temperature

      Until now, we've been considering how to use minerals in fossils to estimate ancient temperatures and past extents of glaciation. We can also use glacial ice, which exists in layers that we can sample as if they were layers of sediment: older lower and younger higher.

      We said that, as water vapor moves toward the poles, it is increasingly enriched in ¹⁶O and depleted of ¹⁸O. The extent to which that happens depends on temperature. In this case, where we're talking about making snow from vapor (not a mineral from solution), colder conditions lead to a lower δ¹⁸O of the snow. Glacial ice thus provides a record of temperature over the poles, but its temperature dependence is opposite the one we talked about for minerals in Part 1.

Summary

      The upshot of all this is that we can use oxygen isotope data from fossils and from glacial ice to unravel how global climate has fluctuated. The generalizations from each of the above sections are:

1. Greaterδ¹⁸O of a mineral results from colder temperatures where that mineral formed.

2. Greater δ¹⁸O of a marine mineral results from more glaciation.

3. Lower δ¹⁸O of glacial ice results from colder polar temperatures.

      Note that Items 1 and 2 work together: If we're examining a fossil from the ocean, a colder world and a glaciated world both cause a greater δ¹⁸O of the mineral material in that fossil. That coincidence has made it a little harder to distinguish specific temperature effects from specific glacial effects, but it means that we have a robust record of the fluctuations between a colder more glaciated world to a warmer less glaciated world over the last two million years.

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