Mixed layer of the ocean

As a result of an accident a cloud of hexachlorobenzene (HCB) of approximatively circular shape floats in the surface-mixed layer of the ocean. At time t0 when the patch is first detected, it has approximately the shape of a two-dimensional normal distribution ... 

If the vertical distribution of particulate thorium is uniform, as would be the case in the surface mixed layer of the ocean for example, and if the system is at steady state, then Equation (3) reduces to... 

The first relaxation time is associated with the equilibration of excess l4C between the atmosphere and the mixed layer of the ocean the second relaxation time describes the much slower rate of filling up the deep sea... 

Consider now a perturbation of the system, specifically the addition of C02 to the atmosphere. Equilibration with the mixed layer of the ocean is achieved quickly within a few years, although the evasion factor prevents a favorable partitioning of C02 between both reservoirs and more than 90% of the added C02 stays initially in the atmosphere. The excess is eventually transferred into the deep sea. Complete equilibrium is achieved slowly and requires a time of the order of 1000 yr. In the interim period, the biosphere may utilize some of the excess C02 and store it temporarily as organic carbon in biomass and soil humus. The residence time is longest in soil humus (see Table 11-10), amounting to some hundred years, on average. This is still short compared with the time needed to attain complete equilibrium between atmosphere and ocean, so that it is the latter process that determines the long-term behavior of the system. 

Here D is the molecular diffusivity of CO2, z is the film thickness, a, is the solubility of i, V and A are the volume and surface area of the ocean, and X is the decay coefficient. Use of pre-industrial mean concentrations gave a global boundary layer thickness of 30pm (D/z 1800m y = piston velocity). The film thickness is then used to estimate gas residence times either in the atmosphere or in the mixed layer of the ocean. For CO2 special consideration must be made for the chemical speciation in the ocean, and for " C02 further modification is necessary for isotopic effects. The equilibration times for CO2 with respect to gas exchange, chemistry, and isotopics are approximately 1 month, 1 year, and 10 years, respectively. 

First I want to review some relevant information regarding the global carbon cycle and the processes that affect atmospheric concentrations of carbon dioxide. There are vast reservoirs of carbon in the system (see Figure 3.1) that can exchange fairly rapidly with the atmosphere, which contains about 750 gigatons (1 gigaton = 10 tons) of carbon (GtC). The terrestrial biosphere and soils contain about 2,000 GtC the mixed layer of the ocean contains about 1,000 GtC and the deep oceans, 38,000 GtC. 

Other limitations on phytoplankton growth are chemical in nature. Nitrogen, in the form of nitrate, nitrite and ammonium ions, forms a basic building material of a plankton s cells. In some species silicon, as silicate, takes on this role. Phosphorus, in the form of phosphate, is in both cell walls and DNA. Iron, in the form of Fe(III) hydroxyl species, is an important trace element. Extensive areas of the mixed layer of the upper ocean have low nitrate and phosphate levels during... 

The forms of equations 14 and 17 describing the deposition of the tracer on the ocean surface, ocean floor and into the historical layers of the sediments are all similar. However, the amplitude variations in the historical layers of sediment are attenuated considerably compared to variations in deposition on the ocean surface, i.e., at input, the attenuation being governed by the effective residence times of nuclides in sea water and in the mixed layer of the sediments. 

Mercury is a naturally occurring element. Natural emissions of mercury, e.g. from ore deposits and from volcanic activity, are variously estimated at amounts between 2500 and 5500 tonnes/year and are thus similar in magnitude to anthropogenic emissions, which are currently estimated at some 3600-4100 tonnes/year world-wide. Some 30000 tonnes of mercury are readily available in the environment, i.e. in the atmosphere or in the mixing zone of the oceans, with tens of millions of tonnes in the upper layers of the continental masses and still more in the deep oceans (see Table 2.1). 

Th, Co, and, in some locations, Fe. Surfece-water enrichments are usually caused by rapid rates of supply to the mixed layer via atmospheric deposition or river runoff. Removal usually occurs through relatively rapid precipitation into or adsorption onto sinking particles. Trace elements controlled by scavenging tend to have short (100 to lOOOy) residence times. Since these residence times are less than the mixing time of the ocean, significant geographic gradients are common. 

By contrast, the gas transfer estimates utilizing Rn measurements assumes steady state between Rn production from radioactive decay of nonvolatile Rd and gas transfer with the atmosphere. This assumption is possible because Rn has a half-life of only 3.8 days, so accumulation and lateral ocean fluxes of Rn is assumed to be minimal. Again, a potential problem is the active, versus inactive layer of the ocean in this case, the mixed layer depth that may change during an experiment. 

Despite the difficulty of interpreting 14C measurements on surface ocean water such measurements are of great interest. The net transport of excess 14C from the atmosphere to the sea depends on the difference between the 14C concentration in atmospheric C02 and that in the carbonate system at the sea surface. The decline in the atmospheric reservoir of excess 14C is therefore controlled by the 14C concentration at the sea surface. This in turn depends upon diffusion and advection into the deep sea. As the levels of excess 14C in the troposphere and the mixed layer of the sea begin to approach each other, mixing from the mixed layer of the sea into the deep sea will be the factor controlling the levels of excess 14C in the atmosphere. 

The mean residence time of carbon in the mixed layer of the sea before transfer into the deep sea is of considerable interest, for as has already been pointed out, the rate of this transfer will eventually govern the levels of excess 14C in the atmosphere. There have been several estimates of this residence time. Craig (29) concluded that it was most probably not more than 10 years, and in one of his calculations he deduced a value of 4 years. Broecker et al. (14) concluded it was 5 years in the Atlantic Ocean and 8 years in the Pacific Ocean. Nydal (45) found that for the North Atlantic it was around 3 years or less. The profiles of Figure 6, and a few others which are not shown, all show a significant penetration of excess 14C below the mixed surface layer, pointing to a short residence time, of the order of 2 years, in the mixed layer of the sea before transfer below the thermocline into the deep sea. Considering the size of the oceans these data are very meager, and no firm conclusions can be drawn from them. However, continued measurements of 14C in the sea should help to establish a firmer estimate of this quantity. 

Atmosphere-ocean carbon exchange is much controlled by physical processes, including mixing of the surface and deep layers of the ocean across the thermo-cline. Biological processes favor the movement of carbon from the surface layer to deep layers and down to bottom sediments. The biological pump functions as a result of phytoplankton photosynthesis. 

Johnson J. E. and Bates T. S. (1996) Sources and sinks of carbon monoxide in the mixed layer of the tropical South Pacific Ocean. Global Biogeochem. Cycles 10, 347-359. 

To further examine the possible role of various processes in the consumption of methylchloroform, pseudo-first-order rate constants were approximately computed for the mixed layer of different oceanic regions (Table VII). 

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