Forams to the rescue

As it happens, Nature has provided a handy way of determining the isotopic composition of past ocean waters, via the good works of single-celled shelly amoeba-like organisms known as foramanifera, nicknamed forams (see Fig. 1.6). These creatures build distinctive calcium carbonate (CaCO3) shells which record the state of the water in which they grew. Because the shells have such diverse and unmistakable shapes, it is easy to recognize and select out the species which live at the the depth level one wishes to investigate. The two principle types of forams are benthic which live near the sea floor, and planktonic which require light and live near the ocean surface. The shells of both types wind up in tidy layers in the marine sediments, allowing one to read the state of the ocean in both depth and time, from a single sediment core. Benthic forams appear in the fossil record as early as 525 million years ago, but their use as paleoclimate indicators has been primarily restricted to the period over which significant portions of seafloor have survived -approximately the past 70 million years.

Forams are not just passive recorders of the isotopic composition of ocean waters, however. As is the case for any chemically distinct pair of reservoirs in contact, the oxygen in foram carbonate is systematically fractionated relative to the isotopic composition of sea water. There is some dispute as to the extent to which this fractionation can be thought of as equilibrium fractionation, but the fractionation factors behave much like inorganic equilibrium fractionation factors and do not differ greatly amongst species. Carbonate prefers the heavier forms of oxygen, and at a temperature of 18C, the 18 O to 16 O of precipitated carbonate is greater than the ratio for the water in which it precipitates by a factor of about 1.03. In other words, carbonate is about 30%oenriched in 18O compared to the water with which it is in equilibrium. As is typically the case for equilibrium fractionation, the degree of enrichment increases as temperature decreases. The change in fractionation with temperature is usually expressed as a paleotemperature equation. Many paleotemperature equations have been given, based on laboratory-cultured organisms, on field observations of recently living forams, and on laboratory measurements of inorganic carbonate precipitation. All give similar results, though the differences are important if one is interested in high accuracy. A general feel for the numbers is adequately given by the following paleotemperature equation, which applies to the benthic foram Uvigerina.

T = 17.97 — 4.0 • (6c(VPDB) — 6w(VSMOW)) (1.8)

where T is the temperature in degrees C at which the foram grew, 6c(V PDB) is the measured 618O of the foram carbonate reported relative to VPDB, and 6w(V SMOW) is the 618O of the water in which the foram grew, reported relative to VSMOW. Both 6 values in the above equation are to be expressed in permil units.

The temperature dependence of foram fractionation is a two-edged sword. On the one hand, the temperature dependence allows forams to be used as paleothermometers. On the other hand, the temperature dependence means it is hard to disentangle ice-volume effects from temperature effects. According to Eq. 1.8, a 618O variation in foram carbonates of 2%ocould represent an temperature change of 8K where the forams grew, or instead a change of 2%oin the water in which the forams grew - corresponding to an ice volume change equivalent to roughly 200m of sea level. The use of benthic forams mitigates this ambiguity to some extent, since the deep ocean temperature is much more uniform than surface temperature. This is so because the deep ocean is filled with waters that are created at the coldest parts of the surface ocean, typically located near the poles. When the climate is in a state having ice at one or both poles, this temperature hovers around the freezing point of sea water, whence the benthic oxygen isotopes primarily reflect ice volume rather than temperature - though the temperature effect is still by no means negligible. A 2K variation in deep ocean temperature, which is not implausible even in icy conditions, leads to about a 0.5%ovariation in the 618O of carbonates, which if attributed instead to sea water composition would translate into an ice volume variation equivalent to about 50m of sea level. At the other extreme, when the climate is in a state without ice at either pole, the isotopic composition of ocean water itself can be considered nearly fixed, and the benthic foram isotopes provide an indication of the polar temperature, regardless of where the sediment core is actually drilled. In this case, a 2%oincrease in benthic foram 618O would indicate an 8K warming of polar temperature.

Benthic forams thus provide a valuable overall indicator of the state of the climate, giving an indication of polar minimum temperature in ice-free climates, and ice volume in icy climates. Since greater ice volume generally goes with a colder climate, and both high ice volume and low temperature make the 618O of foram carbonate more positive, high benthic foram 618O indicates a cold climate while low benthic foram 618O goes along with a relatively warm climate. In icefree climates, planktonic forams can be used to estimate surface temperature, but in icy climates the need to subtract out the ice volume effect makes it hard to get accurate surface temperature estimates by this means.

Forams also preserve other chemical signatures that are useful in reconstructing the past state of the climate system. Notably, they provide a record of the 613C of inorganic carbon of the ocean. Abiologically precipitated sea floor carbonates, or carbonates not associated with individual microfossils, also do this, but the additional depth information available by using benthic vs. planktonic forams provides valuable information about the state of the oceanic carbon cycle. The use of fractionation as a paleothermometer is not limited to isotopes. Notably, magnesium (Mg) substitutes for calcium (Ca) in foram shell carbonates, to an extent that depends on temperature; hence the Mg to Ca ratio in foram shells can be used as a paleothermometer 6. As with oxygen isotopes, the fractionation is relative to the composition of sea water, but the Mg to Ca ratio of ocean water is not affected by formation of glaciers, and hence evolves relatively slowly over

6 Magnesium-calcium paleothermometry can also be used with cores drilled into corals. The growth zone of corals has a distinct depth preference, which has proved particularly useful in estimating ocean surface temperature

Eon Era

Eon Era Period





i 1 i i i i 1 i i i i 1 i i i i 1 i i i i


J_L 1..

3000 2000


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