4.3.1 The carbonate compensation

As discussed in section 2.3.4, the CaCO₃ burial in sediments is ultimately compensated for by the input from rivers. Because weathering and sedimentation rates appear relatively independent, there is no a priori reason why these two processes should be in perfect balance at any particular time. However, any imbalance between them can lead to large variations of the stock of calcium carbonate in the ocean (and thus of the alkalinity, Eq. 2.45) on millennial to multi-millennial timescales . This would imply significant changes in the oceanic pCO₂ (section 2.3.2) and in the atmospheric CO₂ concentration. Such large shifts were not observed, at least over the last tens of million years, because of a stabilising feedback between the oceanic carbon cycle and the underlying sediment referred to as carbonate compensation.

To understand this feedback, it is necessary to analyse the mechanisms controlling the dissolution of CaCO₃. First, let's define the solubility K^CaCO₃ (similarly to the solubility of CO₂ in Eq. 2.43) from the equilibrium relationship for the dissolution of CaCO₃ (Eq. 2.47).

where [CO₃^2-]_sat and [Ca²⁺]_sat are the concentrations when the equilibrium between CaCO₃ and the dissolved ions is achieved, i.e at saturation. If at one location in the ocean the product [CO₃^2-][Ca²⁺] is higher than K^CaCO₃, the water is said to be supersaturated with respect to CaCO₃. If the product is smaller than K^CaCO₃, the water is undersaturated. As the variations of the concentration in Ca²⁺ are much smaller than the variations in the concentration of CO₃^2-, saturation is mainly influenced by [CO₃^2-].

Observations show that the concentration of CO₃^2- decreases with depth (inset of Fig. 4.13). The downward transport of inorganic carbon by the soft tissue pump and the carbonate pump might suggest the opposite. However, we must recall that the alkalinity is mainly influenced by the concentration of bicarbonate and carbonate ions (see the discussion of Eq. 2.45). Neglecting the small contribution of carbonic acid, Alk can be approximated by:

If we also neglect the influence of the carbonic acid on the DIC, we can write:

This leads to

The dissolution of calcium carbonate releases CO₃^2- directly. This is consistent with Eq. 4.20 as the dissolution of 1 mole of CaCO₃ increases Alk by 2 and DIC by 1. By contrast, the remineralisation of organic matter mainly affects the DIC, producing a decrease in [CO₃^2-], according to Eq. 4.20. In the present oceanic conditions, the influence of the soft tissue pump appears to dominate, leading to the observed decrease of CO₃^2- at depth. As the solubility of CaCO₃ increases in the deep ocean, mainly because of its pressure dependence, the upper ocean tends to be supersaturated while the deep ocean is undersaturated. The depth at which those two regions are separated is called the saturation horizon (inset of Fig. 4.13).

Some of the CaCO₃ that leaves the surface layer is dissolved in the ocean water column, but a significant part is transferred to the sediment. There, a fraction of the CaCO₃ is dissolved and reinjected into the ocean, the rest being buried in the sediment on a long timescale. In order to describe those processes, the lysocline is defined as the depth up to which CaCO₃ in sediments is subject to very little dissolution while below the Calcite Compensation Depth (CCD) nearly all the calcite is lost from present in the sediment is lost by dissolution. The CCD is then the depth at which the input from sedimentation exactly balances the loss from dissolving calcite at the top of the sediment. The region between the lysocline and the CCD is called the transition zone (Fig. 4.13).

The position of the transition zone depends on several factors, in particular the presence of organic material in the sediment. It is significantly influenced by the saturation of the waters above the sediment: if they are undersaturated, dissolution in the sediment tends to be relatively high, but if the waters are supersaturated dissolution is very low (as in the upper ocean). If the saturation horizon changes, the position of the transition zone is modified and the regions of the ocean where CaCO₃ is preserved in sediment change.

Those shifts in the saturation horizon are responsible for the stabilisation of the ocean alkalinity. Imagine for instance that the input of CaCO₃ from the rivers doubles because of more intense weathering on continents (Fig. 4.13). The alkalinity of the ocean will increase. As a consequence the [CO₃^2-] will increase (Eq. 4.20) and the saturation horizon will fall. The fraction of the ocean floor which is in contact with undersaturated waters will increase. This would lead to a higher accumulation of CaCO₃ in sediment and thus a larger net flux of CaCO₃ from the ocean to the sediment. This fall in the saturation horizon will continue until a new balance is achieved between the increased input of CaCO₃ from the rivers and the greater accumulation in the sediments.

This feedback strongly limits the amplitude of the variations in alkalinity in the ocean and thus of the atmospheric CO₂ concentration. For instance, it has been estimated that, in the example presented above, a doubling of the river input of CaCO₃ would only lead to a change of the order of 30 ppm in the concentration of atmospheric CO₂.

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