The equations shown in Figure 1 provide a succinct summary of the overall net effect of weathering over the long periods of time, when carbonate precipitates locking carbon into a mineral form. However, these equations do not capture the complete effect over shorter time scales, in which dissolved cations from weathering contribute to the total alkalinity (TA) [Dickson, 1981; Wolf-Gladrow et al., 2007] of the oceans and not all cation charge supplied by weathering is balanced by increased oceanic HCO 3 – (as illustrated in the simplified equations in Figure 1). The following equations describe total alkalinity (TA) and dissolved inorganic carbon (DIC) of the oceans [Zeebe and Wolf-Gladrow, 2001]
The equilibrium constants are functions of temperature, salinity, and pressure and thus differ between seawater and freshwater. The whole carbonate system shown above works in concert to determine the relative proportions of the different species of DIC. For present-day sea surface conditions, the relative molar distribution of DIC into its three species H2CO3 * , HCO 3 – , and CO 3 2- is about 1:90:9. Note [H 2 CO 3 * ] = [CO 2 ] + [H 2 CO 3 ]. Variations in these proportions can significantly alter the effect of weathering- derived alkalinity on the amount of CO 2 uptake from the atmosphere. [ 27 ] Let us consider this in the case of Mg-olivine, forsterite (referred to as olivine in the following). This mineral dissolves in water according to the following reaction:
This equation seems to indicate that 4 mol of CO 2 are sequestered during the dissolution of 1 mol of olivine, equivalent to 1.25g CO 2 (or 0.34g C) per g olivine (the molar weight of pure Mg-olivine is 140gmol ?1 ). However, carbonate system chemistry makes the impact of Mg-olivine dissolution on the carbon cycle more complicated than suggested by equation (8), because both DIC and TA are changed, leading to a new, lower, steady state CO 2 concentration. Thus, the ratio of CO 2 sequestration to olivine dissolution will vary with the initial state of the ocean water and with the amount of olivine dissolved. The value of 1.25g CO 2 per g Mg-olivine represents an upper theoretical limit based on the stoichiometry of equation (8). Seawater, assumed to be initially in equilibrium with the atmosphere, will become undersaturated with respect to CO 2 by addition of TA from weathering and will slowly (over weeks to months) reequilibrate by taking up atmospheric CO2 . The amount of CO2 taken up by the ocean is a nonlinear function of initial TA, pCO 2 (atm), temperature, and salinity [Zeebe and Wolf-Gladrow, 2001]. For large amounts of olivine, it is also a function of the amount of TA added. This makes the system seem to some extent complicated, although the calculation is straightforward for a given initial seawater composition and a given addition of alkalinity from weathering. Typical ratios of CO 2 consumption as a function of the amount of olivine-derived alkalinity added to the global oceans and for different starting atmospheric pCO 2 are shown in Figure 5. In general, for the ranges modeled here, the efficiency of carbon sequestration is significantly lower than the theoretical limit of 1.25g CO 2 per gram of Mg-olivine.
The surface ocean is supersaturated with respect to some carbonate minerals. Given this, the input of additional alkalinity from Enhanced Weathering might be expected to promote carbonate precipitation (see the right-hand side of the carbonate equation in Figure 1), which would reduce or reverse the effectiveness of Enhanced Weathering since the carbonate precipitation reaction drives CO 2 release to the atmosphere. However, the abiotic rate of carbonate precipitation is limited in the surface ocean by the presence of sulfate (SO 4 2? ) and phosphate (PO 4 3? ) anions (Mg 2+ cations also inhibit calcite precipitation) [Berner, 1975; Morse et al., 1997; Morse et al., 2007]. The limit to which the marine carbonate system can be modified before driving appreciable rates of carbonate precipitation is not fully understood but is potentially large when distributed globally. Nonetheless, it is necessary to quantify the exact saturation limit for various local surface ocean conditions at which abiotic and biotic precipitation of carbonates would occur.
Excerpted from page 119 paragraph 26 of this following paper, with special thanks to Jens Hartmann, A. Joshua West, Phil Renforth, Peter Köhler, Christina L. De La Rocha, Dieter A. Wolf-Gladrow, Hans H. Dürr, and Jürgen Scheffran for their work in this area: