Geologic carbon sinks reaction with sedimentary carbonates

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Where does the 'geologic' part come into the picture? In shallow water environments, carbonates are precipitated by corals and ben-thic shelly animals, with a smaller 'abiotic' contribution occurring as fill-in cements and coatings on mineral grains and biogenic mat ter. Approximately 0.3 Pg C equivalent of calcium carbonate (CaCO3) is produced annually in these environments (Milliman and Droxler, 1996). Carbonate is also precipitated biologically in the open ocean (i.e. away from the continental shelf) by plankton such as coc-colithophores and foraminifera, as well as by pteropods. About 0.8 Pg C/year of CaCO3 is thought to be produced here (Milliman and Droxler, 1996; Feely et al., 2004). What happens to this carbonate?

The shallow waters of the ocean margins are everywhere oversaturated with respect to the solid CaCO3 phase (i.e. the saturation state, W > 1.0; see Box 6.1). The dissolution loss of carbonate is therefore relatively small; of the ~0.3 Pg C/year produced, about 0.17 Pg C/year is thought to be buried virtually in situ while another 0.04 Pg C/year is exported to the adjoining continental slopes (Milliman, 1993; Milliman and Droxler, 1996). The total neritic accumulation of CaCO3 today is therefore ~0.2 Pg C/year, and thousands of years of build-up of this material has given rise to large-scale topographical features such as barrier reefs and carbonate banks and platforms. The shallow-water accumulation rate encapsulated in these estimates is probably significantly higher than the long-term (glacial-interglacial, or >100,000 years) average because reef growth rates are still adjusting to the rise in sea level that accompanied the termination of the last glacial period (e.g. Ryan et al., 2001; Vecsei and Berger, 2004).

The situation is quite different in the open ocean because oceanic waters become increasingly less saturated at greater depth (and increased pressure). When the ambient environment becomes undersaturated (W < 1.0) carbonates will start to dissolve. This occurs at ~4500 m in the Atlantic Ocean and ~3000 m in the Pacific Ocean. At more than 1000 m deeper than this, sediments are typically completely devoid of any carbonate particles. Topographic 'highs' on the ocean floor such as the mid-Atlantic ridge (where the ocean floor is 'only' ~3000 m deep and W > 1.0) can thus be picked out by sediments rich in CaCO3 while the adjacent deep basins (W < 1.0) are low in CaCO3 content (Fig. 6.6); an effect likened to 'snow-capped

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Fig. 6.6. Observed distribution of the calcium carbonate content (as percentage dry weight (wt%) in each sample) of the surface sediments of the deep sea. (From Archer, 1996a.) Areas with no data coverage (parts of the Southern Ocean, and many of the continental margins) are left blank.

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Fig. 6.6. Observed distribution of the calcium carbonate content (as percentage dry weight (wt%) in each sample) of the surface sediments of the deep sea. (From Archer, 1996a.) Areas with no data coverage (parts of the Southern Ocean, and many of the continental margins) are left blank.

mountains'. The situation is actually much more complicated than this and other factors are required to completely explain the whole observed pattern of sediment composition, such as the breakdown of organic matter in surface sediments and the release of metabolic CO2 into the sediment pore-waters (see Archer, 1996b; and Ridgwell and Zeebe, 2005). At a global scale, only ~10—15% (equal to 0.1 Pg C/year) of carbonate produced at the surface ever escapes dissolution to be buried in accumulating deep-sea sediments (Milliman and Droxler, 1996; Archer, 1996b; Feely et al., 2004).

Although CaCO3 burial in shallow waters currently appears to be greater than in deep-water sediments, we will focus on the latter carbon sink. We justify this simplification because carbonate preservation and burial in the deep sea is much more important in regulating atmospheric CO2 (Archer, 2003). Although it is beyond the scope of this chapter, note that sea-level and evolutionary changes occurring during the Phanerozoic period have altered the balance between shallow- and deep-water CaCO3 burial. This must have had a profound impact on the regulation of ocean chemistry and climate (Ridgwell et al., 2003; Ridgwell, 2005).

Carbonate burial (in the deep ocean) represents a 'geologic' sink of carbon. How does this relate to the long-term fate of fossil fuel CO2? The first mechanism we will discuss is illustrated in Fig. 6.7a. Dissolution of CO2 in surface waters results in a decrease in ambient carbonate ion (CO2-) concentration (see Figs 6.4c, 6.7c). Think about shifting the aqueous carbonate equilibrium reaction CO2(aq) + CO|- + H2O « 2HCO- to the right to (partly) compensate for the addition of CO2. Because a reduction in COf- reduces the stability of carbonates (see Box 6.1), the invasion of fossil fuel CO2-enriched waters into the deep ocean will drive an increase in the rate of dissolution of CaCO3 in the sediments (Sundquist, 1990; Archer et al., 1997, 1998). If this rate of dissolution overtakes the rate of supply of new biogenic carbonate from above, previously deposited carbonate will start to dissolve (erode) (Fig. 6.8a).

Each mole of CaCO3 that dissolves removes one mole of CO2(aq) to form two moles of bicarbonate:

(the difference between dotted and dashed HCO3- inventory trajectories in Fig. 6.8c). Thus, water masses that have passed over carbonate-rich sediments become, in a sense, 'recharged', and are able to absorb more CO2 from the atmosphere (Fig. 6.7a). One can think of this as anthropogenic CO2 being 'neutralized' by the reaction with sedimentary CaCO3. We will refer to this carbon

As CaCO, in deep-sea sediments dissolves, C02M is transformed into bicarbonate, in effect rendering it no longer available for exchange with the atmosphere:

An imbalance Is Induced between Inputs to the ocean from (mainly carbonate rock) weathering and carbonate burial losses. Because the carbonate weathering reaction consumes C02 (CO,w + HzO + CaCO, -> Ca* + 2HCO;) on a tlmescale of 10* years, fossil fuel COa Is further removed from the atmosphere and locked up In the ocean.

Ultimately, on a timescale of 10B-10a years, greenhouse-enhanced rates of silicate rock weathering (2COw) + HsO + CaSiOj

-> Ca" + 2HCO3 + Si02) result In the permanent removal of C02 and transfer to the geologic reservoir.

An imbalance Is Induced between Inputs to the ocean from (mainly carbonate rock) weathering and carbonate burial losses. Because the carbonate weathering reaction consumes C02 (CO,w + HzO + CaCO, -> Ca* + 2HCO;) on a tlmescale of 10* years, fossil fuel COa Is further removed from the atmosphere and locked up In the ocean.

Ultimately, on a timescale of 10B-10a years, greenhouse-enhanced rates of silicate rock weathering (2COw) + HsO + CaSiOj

-> Ca" + 2HCO3 + Si02) result In the permanent removal of C02 and transfer to the geologic reservoir.

Fig. 6.7. Mechanisms of carbon sequestration (II). Panels (a) through (c) illustrate the pathways of carbon uptake occurring on timescales of millennia (103 years) and beyond - the 'geologic' carbon sinks.

(a) Operation of the sea-floor CaCO3 neutralization;

(b) operation of terrestrial CaCO3 neutralization; and

(c) operation of the silicate weathering carbon sink.

carbon (DIC) that is in the form of CO2(aq) actually decreases (compare dotted and dashed lines in Fig. 6.8c) as a result of reaction with CaCO3. It is the associated reduction in ambient pCO2 that allows further transfer of CO2 from the atmosphere to the ocean.

Clearly we need to quantify how much CaCO3 will dissolve from the sediments and what effect it will have on the removal of CO2 from the atmosphere. We therefore use an extended carbon cycle model that includes the relevant interaction with carbonates in deep-sea sediments (Box 6.2). The evolution of atmospheric CO2 is shown in Fig. 6.8b. Now, CO2 is declining slightly faster at year 3000 compared with when the ocean invasion sink is operating alone. It is important to recognize, however, that the peak atmospheric CO2 value reached (and thus the maximum extent of 'global warming') is virtually unaffected by the inclusion of the buffering of ocean chemistry by carbonate-rich sediments. The effect of sediment dissolution is rather more pronounced over the following few thousand years, and atmospheric CO2 reaches a new, lower steady state of 715 ppm not long after year 10,000. Thus, neutralization with sea-floor carbonates eventually results in the additional removal of 444 Pg C from the atmosphere, or about 11% of the initial fossil fuel burn that we assumed. Again, the relative importance and fraction of CO2 sequestered by this mechanism depends on the magnitude of the fossil fuel burn - sea-floor carbonate neutralization has been estimated to account for 9.0% of an 874 Pg C fossil fuel release, rising to 14.8% for a 4550 Pg C release (Archer et al., 1998). Thus, the buffering response in the GENIE-1 model is slightly less than in a previous model study.

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