By Dr. Ian Walsh, Consulting Chief Scientist @ Ocean-based Climate Solutions, Inc

In the ocean we have challenges to measuring the impact of ocean upwelling and deriving time weighted sequestration fluxes.

But these are entirely tractable challenges: we have tools and experience through generations of oceanographers in measuring carbon flux through the ocean where we can and estimating where we don’t have data. We have local process studies and global models that provide bounds to carbon fluxes from diel to annual and greater time scales. Global models typically use well structured physical models (get the physics right first!) overlayed with biogeochemical models with more or less stocks and pathways to apportion energy and variability over significant time scales.

The challenge with ocean upwelling is that along with nutrients, the dissolved inorganic carbon (DIC) concentration also has a regeneration profile, with low concentrations at the surface relatively monotonically increasing with depth, with the gradient controlled mostly by water mass.

So we have to measure both the nutrients and the carbonate chemistry of the upwelled water. What we want with the upwelled water source is a source water for upwelling where there is an ‘excess’ of nutrients relative to the DIC such that there is a net uptake and fixing of atmoCO2 into the particle pool via phytoplankton rather than simply a quick shunting of upwelled CO2 as DIC into the fixed carbon pool. The presence of ‘excess’ nutrients, particularly phosphate, relative to DIC in the twilight zone is a steady state result of nutrients having higher first order remineralization rates than particulate organic carbon. This dynamic operates over long periods of time and space driving the linkage of atmoCO2 to ocean upwelling on a global scale through recent geological time has been long established (Dymond and Lyle, 1985)

There are many ways to define the ‘excess’ relative nutrient load relative to the carbon load. Karl and Letelier (2008) used the phosphate concentration relative to nitrate and DIC to define a sequestration potential for a given depth profile. The attraction for using this method is that temporal and spatial variability of dissolved constituents in the ocean decreases with depth and particularly in the central gyres a relative sparsity of data still yields volumetric/depth/time relationships that are reasonably stable, and hence the initial conditions of the setting of a particular pump deployment in a central gyre can be reasonably assumed to be the conditions over annual time scales.

The task then is to gauge the actual delivered seawater. For that we can use a mixing model over the volume addition:

Temperaturez,plume = A * Temperaturez,ambient + D * Temperatureintake

PO4ex plume = A * PO4z,ambient + D * PO4intake


A + D = 1

Since for the oligotrophic ocean we can assume that the available dissolved nutrient concentration in the mixed layer is non detectable, i.e.:

PO4ML,ambient = 0

Where z here is replaced by the mixed layer depth.

If we assume that the upwelled volume is dispersed into the plume, and that the plume is entirely within the mixed layer, and the mixed layer is equal to or less than the euphotic depth, then we can presume that all upwelled water that enters at the intake is dispersed within the mixed layer and is taken up over time into the biological system.

Since we know that the ambient excess phosphate is zero, then the additional phosphate added to the system will also devolve to zero and that this will occur on the order of days. Hence, the simplest measurement of the net sequestration effect of a given artificial upwelling pump can be based on the temperature and nutrient load relationship at the intake depth range and the measured temperature anomaly at the surface.