This is because oceanic waters become increasingly less saturated with depth. Below the depth of the saturation horizon, conditions become under-saturated (X b 1.0) and carbonate will start to dissolve. In the modern ocean the calcite saturation horizon (see Box 1) lies at about 4500 m in the Atlantic and ~ 3000 m in the Pacific Ocean. Within a further 1000 m sediments are often completely devoid of carbonate particles (the carbonate compensation depth, or dCCDT). Topographic highs on the ocean floor such as the midAtlantic ridge can thus be picked out by sediments rich in CaCO3 while the adjacent deep basins are low in %CaCO3 (Fig. 2). The visual effect has been likened to dsnow-capped mountainsT. The pressure induced surface-to-deep vertical contrast in X is further enhanced by the respiration of organic matter and release of metabolic CO2 in the ocean interior which suppresses the ambient carbonate ion concentration and thus X (see Box 1). The greater accumulation of metabolic CO2 in the older water masses of the deep Pacific explains why the sea-floor there is much poorer in %CaCO3 compared to the Atlantic at a similar depth [4] (Fig. 2). Unfortunately, the carbonate cycle does not conform to this simple picture, and a significant fraction of CaCO3 appears to dissolve in the water column even before it can reach the sediment surface [4–6]. This has been something of an enigma because the reduction in carbonate flux measured by sediment traps occurs well above the depth at which calcite becomes thermodynamically unstable. Dissolution of carbonate particles in acidic digestive conditions of zooplankton guts has been one proposed mechanism [6]. Acidic micro-environments within individual dmarine snowT aggregates may also be important [7]. Another possible explanation surrounds the A. Ridgwell, R.E. Zeebe / Earth and Planetary Science Letters 234 (2005) 299–315 303 90 100 60 0 40 wt% CaCO3 80 20 0 -90 0 90E 180E 270E 0 Fig. 2. Distribution of the calcium carbonate content of the surface sediments of the deep sea [10]. There is an apparent predominance of CaCO3 accumulation taking place in the Atlantic and Indian Oceans compared to much more sparse accumulation in the Pacific. This is primarily a consequence of the greater accumulation of metabolic CO2 in deep Pacific waters which drives a greater degree of under-saturation and lowers the depth of the lysocline (see Box 1). The virtual absence of CaCO3 in sediments of the Southern Ocean is due to a combination of much lower CaCO3/POC rain ratio to the sediments and relatively corrosive bottom-waters. Topographic dhighsT can be picked out as areas of higher wt.% CaCO3 compared to sediments elsewhere in the same basin at similar latitudes. Areas with no data coverage (parts of the Southern Ocean, and many of the continental margins) are left blank. aragonite polymorph because it becomes susceptible to dissolution at much shallower depths than calcite under the same ambient conditions. In support of this are recent estimates of the depth at which most CaCO3 dissolution occurs in the ocean which appears to correspond to the aragonite saturation horizon [4,5]. However, calculations suggest that solute release from sinking pteropod shells, the main aragonite product in the open ocean, should mostly occur much deeper than this [8]. Dissolution of aragonite also does not help explain how 65% of calcitic foraminiferal tests can be lost at shallow depths [9]. This uncertainty is of concern because a full appreciation of the controls of atmospheric CO2 and response to global change requires an understanding of the dissolution and the depth of recycling of CaCO3 in the water column. Overall, more than 80% of all carbonate precipitated in the open ocean dissolves either in the water column or within the uppermost layers of the underlying sediments [4,10,11]. The remainder, some 1 Gt CaCO3 yr 1 accumulates in deep-sea sediments. This burial loss, to which can be added as much as another 1 Gt CaCO3 yr 1 of deposition in neritic environments (although the uncertainty in this figure is substantial) [11,12], must somehow be balanced if the ocean is not to run out of calcium ions! This is achieved through the weathering of carbonate and silicate rocks. 1.2. Weathering and carbonate recycling The weathering of calcium carbonate and calcium silicate minerals in soils and at exposed rock surfaces helps balance the CaCO3 sedimentation loss by unlocking Ca2+ from the geologic reservoir. Alteration of ocean crust by percolating fluids adds an additional but more minor contribution [13]. The weathering reactions (see Fig. 1b; #5 and #7) provide the other raw material necessary for carbonate precipitation— bicarbonate ions (HCO3 ). However, because the transformation 2CO2 Y 2HCO3 (Fig. 1b; #7) is internal to the surficial system and does not represent a source of dnewT carbon, the component of CaCO3 burial derived from silicate rock weathering represents a loss of carbon to the geologic reservoir. This must be replaced on the long-term, achieved through the release of CO2 to the atmosphere from volcanic sources2 [1]. (In contrast, the weathering and burial of CaCO3 results in no net loss or gain of CO2 to the surficial system). 2 Imbalances between the rates of burial of organic carbon and weathering of ancient organic matter (kerogens) exposed at the land surface affects the inventory of carbon in the surficial reservoirs and thus atmospheric CO

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