Productivity

Phytoplankton need light, CO2, O2 and nutrients to grow and reproduce. The nutrients

they need are a source of nitrogen – nitrate, nitrite, ammonium or (for nitrogen fixers) dissolved nitrogen; phosphate; orthosilicic acid, hereinafter silicate, (for some classes, including diatoms) and micronutrients, such as iron, zinc and manganese. Any of the above could be a limiting factor, although generally it is not magnesium, calcium, potassium, sodium, sulphate, chloride, CO2 or O2 in the open ocean (Lalli and Parsons

1997).

Light is clearly a limiting factor in winter at high latitudes and can be limiting at other times if the mixed layer is deeper than the critical depth. This is defined as the depth over which depth averaged losses (respiration and grazing) are balanced by depth averaged growth through photosynthesis. There needs to be sufficient forcing for the mixed layer to frequently overturn for this approach to be valid (Huisman et al. 1999). In some cases, often before restratification in spring, the mixed layer is quiescent enough for there to be little transfer of phytoplankton between the euphotic zone (zone with enough

photosynthetically available radiation for net growth of phytoplankton at a fixed depth) and the part of the water column between the euphotic zone and the base of the mixed

layer. This can lead to phytoplankton growth occurring before density stratification. In areas where the mixed layer is very variable there will also be losses through dilution as the mixed layer deepens and some phytoplankton will then remain below the temporary pycnocline once restratification occurs.

1.3.2 Carbon export

The uptake of carbon by phytoplankton and the subsequent sinking of the carbon to deep waters and the sediment (possibly via zooplankton grazing) removes CO2 from the

surface waters of the ocean. Surface waters equilibrate with the atmosphere and so this is a sink of atmospheric CO2 (Figure 1.3). This has been termed the biological pump

(Eppley 1972). The timescale of overturning of the world’s oceans is 1500 years on average so this can represent a long-term removal of carbon from the atmosphere, relative to the time scales of anthropogenic influence on CO2 concentrations. The small

percentage of carbon that enters the sediments is removed for a much greater period, probably >107 years. Under constant conditions, the removal of carbon to deep waters is balanced by outgassing of CO2 from upwelled waters, but an increase in the biological

pump does represent a long-term removal of carbon so long as the increase is away from a strong upwelling region.

Artificially increasing the biological pump has been proposed as a method for sequestering anthropogenic CO2 and studying the efficiency of such a method is one

context of the CROZEX project. Measuring the effects of iron enrichment from ships on deep carbon export has proved difficult, possibly due to the relatively small scales of the blooms created and the timescales of observations.

It has also been hypothesised that the biological pump was stronger in the Southern Ocean during glacial periods (Martin 1990) due to increased dust input (that being due to reduced terrestrial vegetation) although the balance between this and the reduced

effective area of the Southern Ocean (due to increased sea ice extent) means that it is likely that this process can contribute at most 30 ppm of the 80-100ppm change in atmospheric CO2 concentrations observed (Bopp et al. 2003).

Figure 1.3 Schematic of the biological carbon pump (left) and dissolution pump (right).

CO2 is also removed to deep water through the physical dissolution pump (right hand

panel of Figure 1.3). This effect is due to the increased CO2 concentrations of the sinking

water, which has been in contact with present day atmospheric concentrations, relative to the CO2 concentrations in upwelling water, which was mostly exposed to pre-industrial

CO2 concentrations when it formed (Sabine et al. 2004).

1.3.3 Nutrients

The three dimensional flow described in section 1.2.1 brings water from close to the bottom of the Southern Ocean to the surface, leading to a very strong vertical supply of nutrients to the surface layers. Estimating the upwelling from Ekman divergence, Pollard

et al. (2006) estimated the vertical supply of nitrate and silicate to be 13.7 and 50.8 Tmol year-1 respectively. Martin (1990), using an upwelling of 60Sv for the Southern Ocean and an estimate for nitrate concentration and the Redfield ratio, calculated a potential new production of 40 mmol C m-2 day-1 for the summer months. The discrepancy

between this and the observed production of 1.6 mmol C m-2 day-1 is used as support for his hypothesis that iron is limiting in the Southern Ocean.

Figure 1.4 shows that the resulting surface concentrations of nitrate are higher than in any other part of the world’s ocean. The distribution of phosphate, to first order, is the same as for nitrate but at 1/16th of the concentration (the Redfield ratio) so is not shown. Most of the world’s ocean is limited by the availability of macronutrients (nitrate and

phosphate) so this large supply of macronutrients to the Southern Ocean makes it possible that there could be extremely high productivity in the Southern Ocean. There is also a similar pattern to the distribution of silicate in the surface waters (Figure 1.5) although the peak concentrations are further south, linked to the deeper remineralisation of silicate exported from the surface ocean. Silicate is not a prerequisite for

phytoplankton but where it is present siliceous organisms, especially diatoms, often dominate, especially when productivity is high (Hoffman et al. 2006).

Figure 1.4 Climatological surface nitrate concentrations (μmol/l) from the 2005 World Ocean Atlas.

Figure 1.5 Climatological surface silicate concentrations (μmol/l) from the 2005 World Ocean Atlas.

The distribution and biological utilisation of silicate is such that, unlike nitrate and phosphate, it reduces to limiting values over parts of the study area during the growing season. For this reason changes in silicate are sometimes considered in this work but changes in nitrate and phosphate are not. Although silicate is not itself necessary for phytoplankton in general, a reduction of silicate to limiting concentrations when the phytoplankton community structure has developed in the presence of silicate can have significant consequences. It will force a faunal shift towards non-siliceous species that have until that point been out-competed by the diatoms.