CO2aq CO2g H2Ol 2HCO3aqeqn

This reaction makes seawater about eight times more effective at absorbing CO2 than a solution of similar ionic strength but not containing CO2-.

The discussion above assumes that equilibrium is achieved between the sea-water and the air with respect to CO2. This leads to the second factor which must be taken into account, since the slow mixing time of the oceans means that it takes hundreds, if not thousands, of years for equilibrium to be attained over the whole depth. In general, it is not transfer across the sea surface which is rate-limiting for uptake of CO2, but mixing of surface water down to the ocean deeps (mean depth 3.8 km, maximum depth 10.9 km). Such mixing is greatly impeded by the existence in most ocean basins of a stable two-layer density structure in the water. At a depth of a few hundred metres there is a region of rapid temper ature decrease, the main thermocline. This results in enhanced stability of the water column, which inhibits mixing from above or below. It is only in some polar regions, particularly around Antarctica and in the Greenland and Norwegian Seas in the North Atlantic, where the absence of the thermocline allows direct, and therefore rapid, mixing of surface with deeper waters (see also Section 6.7).

The large, natural, two-way flow of CO2 across the sea surface makes it very difficult to measure directly the rather small additional flux (about 2% of the gross flux in either direction) resulting from human additions of CO2 to the atmosphere. Best estimates from this approach (which rely on measurements of pCO2 (see Box 3.1) across the sea surface) are about 2 GtCyr-1. In these circumstances resort is often made to mathematical modelling approaches. These models can be of considerable complexity—Box 7.1 shows the principles on which they operate. From modelling studies, the best estimate of the amount of anthropogenic CO2 being taken up by the oceans is 1.9 ± 0.6 GtCyr-1, which is in reasonable agreement with the estimates from direct measurements mentioned above.

As an illustration of how marine biological processes may affect the ability of the oceans to take up CO2 we now briefly discuss the results of some recent field experiments on the role of iron in controlling photosynthesis in the sea (see also Section 6.6). For many years it had been speculated that in some major ocean areas availability of iron was the limiting factor for phytoplankton growth. However, it was only recently that several direct tests of this idea were carried out by adding iron (in the form of ferrous sulphate) to a small (about 100 km2) area of ocean and observing any resulting effects over periods of days. What was found was that addition of only a very small amount of iron led to a dramatic increase in plankton growth (see Fig. 6.26) with resultant drawdown of CO2 from the water (and hence potentially from the atmosphere). This is clear in Fig. 7.4 where pCO2 is lower inside the iron-fertilized patch in relation to the values outside the patch. This is an exciting result in helping to understand what controls marine biological activity. However, its importance for long-term uptake of CO2 by the oceans is not yet established since it may be that much of the extra carbon incorporated into new phytoplankton growth is rapidly recycled by respiration/decomposition in the near-surface waters. Carbon will only be removed from the atmosphere/surface ocean system for any length of time if dead plankton remains sink into the deep ocean. In order to answer this question it will be necessary to carry out iron-fertilization experiments which last for longer and cover a greater area. This will take a considerable effort at the international level and is currently being planned through programmes such as IGBP (Section 7.1).

Fossil fuel burning

It is relatively easy to quantify the amount of CO2 that results from the burning of fossil fuel and other industrial activities, such as the manufacture of cement (as part of this process calcium carbonate (CaCO3) is heated to a high temperature and decomposes, yielding CO2). This source is easier to estimate than those discussed earlier because there is no natural component. All that is required is quan-

Box 7.1 Simple box model for ocean carbon dioxide uptake

In order to calculate how much anthropogenic carbon dioxide (CO2) the oceans can take up from the atmosphere, it is often necessary to construct a model of the system. The simplest of these models divide the oceans into a series of boxes (numbering from a few to several hundred), with water containing its dissolved carbon (C) flowing between them. The main elements of such models are shown in Fig. 1.

For the relatively well-mixed atmosphere and surface-ocean boxes, the carbon flow between them is assumed to be proportional to their carbon content. Within the deep ocean, where the circulation is much more sluggish, vertical mixing is often modelled as a diffusion process. In addition, the model can include a simple circulation with direct input to the ocean bottom from the surface, balanced by upward water movement throughout the deep ocean, to represent convective processes. The spatial and depth distribution of radioactive substances, such as the isotope 14C (see Section 2.8) (produced both by cosmic rays in the atmosphere and from the detonation of nuclear devices in the

Atmosphere

^ First order exchange -fluxes proportional to concentration in upstream reservoir

^ First order exchange -fluxes proportional to concentration in upstream reservoir

Diffusion

Deep ocean A

Vertical diffusion

Vertical circulation

Fig. 1 The main elements of a simple model for ocean CO2 uptake.

(continued)

1950s and 1960s), can then be used to estimate the rates at which CO2 is exchanged between the atmosphere and surface ocean, its diffusion into the deep ocean and its transport by vertical circulation.

For the well-mixed reservoirs, a conservation equation is written in which gain of 14C by inflow to the box (atmosphere or surface ocean) is balanced by the outflow to other boxes plus radioactive decay (see Section 2.8) of the tracer during its time in the reservoir. For the deep ocean, conservation is described by a partial differential advection-diffusion equation. The diffusion coefficient is chosen to best fit the measured depth profile of 14C in the oceans.

Using the model, the uptake of fossil fuel CO2 can be estimated by integrating forward in time from an assumed pre-industrial steady-state value, while adding to the model's atmosphere the estimated year-by-year release of CO2 from fossil-fuel burning. At each time step, the fluxes of carbon between the various boxes are calculated and the carbon contents and concentration profiles changed accordingly. From such models it is calculated that about 35% of anthropogenic CO2 is absorbed by the oceans.

Days after iron added

Fig. 7.4 Surface seawater CO2 concentration during an iron-fertilization experiment. After Watson et al. (2000).

tification of the various fuel types burned every year and a knowledge of the amount of CO2 each produces on combustion. This latter factor, although well known, varies quite a lot between fuels. For example, for each unit of energy produced, coal forms 25% more CO2 than oil and 70% more than natural gas. This occurs because, in the combustion of gas and oil, a major proportion of the energy comes from conversion of hydrogen atoms (H) in the fuel to water (about 60%

Fig. 7.5 Global CO 2 emissions from fossil-fuel burning (solid, liquid and gaseous fuels), cement production and gas flaring for 1751-1999. After Marland et al. (2002).

in the case of gas), rather than from the conversion of carbon to CO2, which provides 80% of the energy when coal is burned.

Recent data on CO2 inputs to the atmosphere from fossil fuels and other anthropogenic sources have been compiled by the United Nations in their Energy Statistics Database. Earlier data have been obtained from a variety of sources but are more uncertain than the numbers for recent years. The results are presented in Figs 7.5-7.7. In Fig. 7.5 the yearly inputs show a general increase over the period since 1751 when records first become available. The data are plotted on a logarithmic CO2 emission scale in Fig. 7.6, which shows that the increase has not always been at the same rate. Although for the periods 1860-1910 and 1950-70 the growth rate was close to 4%, during the two world wars, in the great industrial depression of the 1930s and since the 1970s the rate of increase has been closer to 2%. The slackening of emissions in the last 25 years is due to large increases in the price of oil at the beginning of the period, conservation measures generally and economic retrenchment in the 1990s. Wars, like depressions, are apparently times of reduced economic activity. In Fig. 7.7 the data are plotted by latitude for 1980 and 1989, which clearly show how strongly emissions are skewed towards the industrialized mid-latitudes of the northern hemisphere. Over the 1980s there is a clear shift in emissions southwards, as industrialization has become more global. In the last year for which full data are available (1991), emissions from fossil fuel burning, etc. are estimated to be 6.2 GtCyr-1, with an uncertainty of less than 10%. Average annual emissions over the 1980s were 5.4 ± 0.3 GtCyr-1. Finally, in Fig. 7.8 the record of fossil fuel emissions and atmospheric concentration increase at Mauna Loa from 1958 to 2000 is shown. This illustrates the variability of the increase in atmospheric CO2 year by year (discussed later), and also that only about half of the carbon emitted stays in the atmosphere to produce the yearly increase (as mentioned previously).

1920 1940 Year

Fig. 7.6 Global annual emissions of CO2 from fossil-fuel combustion and cement manufacture. After IPCC (1990). With permission of the Intergovernmental Panel on Climate Change.

1920 1940 Year

Fig. 7.6 Global annual emissions of CO2 from fossil-fuel combustion and cement manufacture. After IPCC (1990). With permission of the Intergovernmental Panel on Climate Change.

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