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.
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.
Back to Railsback's GEOL 1122 main page
Back to Railsback's main page
Back to the UGA Geology Home Page