The most notable aspect of present-day carbonate weathering is that it is so much faster than silicate weathering. Silicate terrains, such as portions of the high Himalayas and the New Zealand Alps (Blum et al., 1998; Jacobson et al., 2003) that contain only traces of carbonates are marked by water chemistries dominated by carbonate dissolution. On a much grander scale, although silicates cover most of the surface area of the continents, the global average chemical composition of river water is characteristic of that expected for carbonate weathering (Meybeck, 1987; Berner and Berner, 1996; Gaillardet et al., 1999). In other words, limestone weathering dominates global water chemistry.
The factors affecting the weathering of Ca and Mg carbonates (cal-cite and dolomite) are similar to those affecting silicate weathering. This includes temperature, as it is affected by the atmospheric greenhouse effect, solar radiation, and continental drift; hydrological changes accompanying changes in global temperature and paleogeography; and the rise and evolution of land plants. However, relief is not as important as it is for silicate weathering. This is because limestones undergo subsurface dissolution by groundwaters regardless of relief. An outstanding example of the development of subsurface limestone dissolution on flat ground is shown by the common sink holes of southern Florida, USA, which can be easily observed from the air by anyone flying into the Miami airport. Also, the flow of water through limestones (and dolostones) is increased as more and more dissolution, including cave formation, proceeds. Silicates require access of water to primary minerals, which can be impeded by a thick, relatively impermeable mantle of clay weath ering products, as in the Amazonian lowlands. Limestone weathering, in contrast, does not involve the production of clays because there is only simple dissolution of the primary minerals.
As pointed out in the Introduction, on a multimillion-year time scale, carbonate weathering has essentially no direct effect on atmospheric CO2. However, it has important indirect effects. First, greater carbonate weathering, for a steady-state ocean, means greater carbonate burial, and greater carbonate burial, if in deep sea sediments, means greater carbonate subduction with greater CO2 degassing. If increased carbonate weathering is forced by higher atmospheric CO2, this leads to positive feedback (Hansen and Wallmann, 2003). Second, in calculating the rate of CO2 uptake by past Ca and Mg silicate weathering, it is necessary to also know the rate of carbonate weathering. Total carbonate buried in sediments comes from the weathering of both carbonates and Ca-Mg silicates. Therefore, to calculate the flux of CO2 taken up by the conversion of Ca-Mg silicates to carbonates (reaction 1.5), one must subtract from total carbonate burial that derived from carbonate weathering (equation 1.13).
Models of the long-term carbon cycle (e.g., Berner and Kothavala, 2001; Wallmann, 2001) include formulation of an expression for carbonate weathering that is similar to that used for silicate weathering. However, in GEOCARB modeling there are some changes from the expression used for silicate weathering: (1) A separate temperature/CO2 feedback factor fBc(CO2) is used that differs from fBt(CO2) because of a different temperature coefficient (Z) for carbonate weathering (Drake and Wigley, 1975); (2) the paleogeographic runoff/weathering term f^t) is not raised to the power 0.65 because carbonates dissolve so fast that they saturate the waters with which they come into contact, and there is no dilution at high runoff (Drake and Wigley, 1975; Stallard, 1995); and (3) the parameter fR(t) is not applied to limestone weathering because of the lesser role of relief, as discussed above.
An additional parameter is added to the expression for carbonate weathering based on the results of Bluth and Kump (1991) for the proportions of land areas underlain by carbonates over the Phanerozoic. Results are derived from a series of global paleolithologic maps laboriously compiled from large amounts of the data of Ronov et al. (1984, 1989). From the results of Bluth and Kump (1991), the dimensionless parameter fLA(t) is derived (Berner, 1994; Wallmann, 2001) that is applied to carbonate weathering:
where fL(t) = fraction of total land area covered by carbonates/the same fraction at present fA(t) = total land area/total present land area.
In the model of Hansen and Wallmann (2003), changes in carbonate weathering fluxes, as a result of changes in the exposure area of carbonates on the continents, are used to drive changes in organic matter burial because the phosphorus released during the weathering of carbonate sediments stimulates organic production upon transfer to the oceans (see figure 3.1).
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