W 33 x 10 9

The difference between the W values in equations 6.9 and 6.11 is mainly due to the effects of ionic strength, which are properly compensated for using total activity coefficients. Overall, the data show that surface seawater is about six to seven times supersaturated with respect to calcite. We might then reasonably expect CaCO3 to precipitate spontaneously in the surface waters of the oceans.

The evidence from field studies is somewhat contrary to the predictions based on equilibrium chemistry. Abiological precipitation of CaCO3 seems to be very limited, restricted to geographically and geochemically unusual conditions. The reasons why carbonate minerals are reluctant to precipitate from surface sea-water are still poorly understood, but probably include inhibiting effects of other dissolved ions and compounds. Even where abiological precipitation is suspected—for example, the famous ooid shoals and whitings of the Bahamas (Box 6.5) —it is often difficult to discount the effects of microbial involvement in the precipitation process.

Volumetrically, the biological removal of Ca2+ and HCO- ions, built into the skeletons of organisms, is much more important. In the modern oceans, the large continental shelf areas created by sealevel rise in the last 11 000 years probably account for about 45-50% of global carbonate deposition. Moreover, about half of this sink for Ca2+ and HCO- ions occurs in the massive coral reefs of tropical and subtropical oceans (e.g. the Australian Great Barrier Reef). It is tempting to assume that coral reefs have always represented a major removal process for Ca2+ and HCO- ions. However, over the last 150 million years it can be shown that it is carbonate sedimentation in the deep oceans which has been volumetrically more important, accounting for between 65 and 80% of the global CaCO3 inventory. These deep-sea deposits, which average about 0.5 km in thickness, mantle about half the area of the deep ocean floor (Fig. 6.9). The ultra fine-grained calcium carbonate muds (often referred to as oozes) are composed of phyto-plankton (coccolithophores) and zooplankton (foraminifera) skeletons (Fig. 6.10). Although these pelagic organisms live in the ocean surface waters, after death their skeletons sink through the water column, either directly or within the faecal pellets of zooplankton.

The controls on the distribution of pelagic oozes are partly related to the availability of nutrients, which must be capable of sustaining significant populations of phytoplankton (see Section 5.5). More important, however, is the dissolution of CaCO3 as particles sink into ocean deep waters. In the deep ocean, carbon dioxide (CO2) concentrations increase, particularly in the deep Pacific, as a result of the decomposition of sedimenting organic matter. Decreased temperature and increased pressure also promote dissolution of CaCO3, favouring the reverse reaction in equation 6.4.

By mapping the depth at which carbonate sediments exist on the floors of the oceans, it is possible to identify the level where the rate of supply of biogenic CaCO3 is balanced by the rate of solution. This depth, known as the calcite compensation depth (CCD), is variable in the world's oceans, depending on the

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