Isotopic Methods to Estimate Marine Productivity

We have two projects that are focused on using isotopic measurements to estimate the rate of biological carbon production and export from the photic layer of the ocean. The first one uses ¹³C/¹²C of dissolved inorganic carbon (DIC) to estimate net export rates of photosynthetically produced organic carbon. The second project uses measurements of the ¹⁸O/¹⁶O and ¹⁷O/¹⁶O of dissolved oxygen to estimate gross photosynthesis rates in the surface ocean. In this discussion δ¹³C is often used to represent the ¹³C/¹²C, where δ¹³C (‰) = [(¹³C/¹²C)x / (¹³C/¹²C)std –1]*1000 and x represents the sample and std represents the standard (PDB).

Using ¹³C/¹²C of DIC to Estimate Biological Carbon Export Rates

The ¹³C/¹²C of DIC in the photic layer primarily depends on three processes: biological organic carbon production, upwelling and mixing of waters from below and air-sea CO₂ gas exchange. Upwelling and mixing at the base of the mixed layer lower the ¹³C/¹²C of the DIC. In contrast, net biological production of organic carbon during photosynthesis increases the ¹³C/¹²C of the DIC. (Biological formation of CaCO₃, because there is little fractionation during this process, has little effect on the ¹³C/¹²C of the DIC.) Air-sea CO₂ gas exchange drives the ¹³C/¹²C toward atmospheric equilibrium. Air-sea CO₂ exchange tends to lower the ¹³C/¹²C in the subtropical ocean and raise the ¹³C/¹²C in the polar oceans (Fig. 1). In the subtropical N. Pacific, the ¹³C/¹²C of DIC is much higher (by up to 2 ‰) than expected at air-sea equilibrium (Fig. 1) and is maintained by the export of organic carbon out of the surface photic layer.

(Fig. 1)

At the times series station ALOHA (23°N 158°W), we have measured the ¹³C/¹²C and concentration of DIC at monthly intervals over the last six years. These data give us a clear picture of the seasonal cycle of these parameters (Fig. 2). Coupling the time rate of change of the ¹³C/¹²C and concentration of DIC allows us to make two DIC budgets for the mixed layer, i.e., one for DIC and one for DIC¹³. Basically, there are three terms in these two budgets, one each for air-sea CO₂ gas exchange, net biological export and upward DIC mixing from below. If we estimate the air-sea CO₂ gas transfer rate based on wind speed measurements and measure the partial pressure of CO₂ gas in the surface water, we can estimate the net air-sea CO₂ flux. Once this CO₂ flux is determined, we use the DIC and DIC¹³ budgets to solve for the remaining two terms, that is, the biological carbon export and upward DIC flux. (We also account for horizontal advection of DIC and DIC13 using measured horizontal gradients and estimates of geostrophic and Ekman transport.) At station ALOHA, the net rate of air-sea CO₂ gas exchange, biological carbon export and upward DIC flux required to explain the summertime DIC decrease and ¹³C/¹²C increase (Fig. 2) are 0.2, 7.2, and 4.1 mmoles m^-2 d^-1, respectively.

(Fig. 2)

Using ¹⁷O/¹⁶O and ¹⁸O/¹⁶O of Dissolved Oxygen Gas to Estimate Biological Photosynthesis Rates

The oxygen isotope method makes use of a naturally occurring ocean-wide isotope labeling experiment to determine both gross and net primary production rates. The method depends on the observation that tropospheric O₂ has an anomalously low ¹⁷O/¹⁶O relative to that expected based on its measured ¹⁸O/¹⁶O. Typically, most natural processes fractionation isotopes in a mass-dependent manner, i.e., the fractionation effect for ¹⁷O/¹⁶O is expected to be about half (0.521) the fractionation for ¹⁸O/¹⁶O because ¹⁷O is one mass unit greater than the common oxygen isotope (¹⁶O) and ¹⁸O is two mass units greater. The ¹⁷O/¹⁶O anomaly for tropospheric oxygen is the result of a mass-independent fractionation that occurs in the stratosphere during reactions between O₂, O₃ and CO₂ induced by UV radiation. These reactions cause stratospheric CO₂ to have a higher ¹⁷O/¹⁶O, and stratospheric O₂ to have a lower ¹⁷O/¹⁶O, than expected if these reactions had fractionation effects that were mass-dependent. Stratosphere-troposphere mixing brings this anomalous ¹⁷O/¹⁶O tagged O₂ into the troposphere.

The ¹⁷O/¹⁶O deviation or anomaly from that expected via mass-dependent fractionations is defined as D¹⁷O = 1000*(d¹⁷O –0.521*d¹⁸O), where d¹⁷O and d¹⁸O represent the ¹⁷O/¹⁶O and ¹⁸O/¹⁶O in per mil (‰) (Luz et al., 1999). Since this mass-independent ¹⁷O/¹⁶O anomaly is small, the units for D¹⁷O are parts per million or per meg (i.e., 1000* ‰). By convention, the D¹⁷O of O₂ in tropospheric air is defined to have a D¹⁷O equal to zero (Luz et al., 1999).

Photosynthesis produces O₂ with a mass-dependent ¹⁷O/¹⁶O and ¹⁸O/¹⁶O. As the amount photosynthetically produced O₂ increases, the D¹⁷O value increases since a D¹⁷O=0 (by definition) represents O₂ in air with the anomalously low ¹⁷O/¹⁶O. If the O₂ was entirely produced via photosynthesis it would have a D¹⁷O=249 per meg. Luz et al. (2000) determined this value experimentally by measuring the D¹⁷O of O₂ produced by marine plankton and corals. Thus oceanic surface waters will have a D¹⁷O somewhere between 0 and 249 per meg depending on the relative rates of air-sea O₂ gas exchange (which drives D¹⁷O towards 0 per meg) and gross photosynthetic O₂ production (which drives D¹⁷O towards 249 per meg). Thus if the rate of air-sea O₂ gas exchange is estimated, e.g., from wind speed measurements, then the rate of gross primary productivity (PPg) is determined from measurements of D¹⁷O. Luz and Barkan (2000) showed that for a mixed layer where biological production (and consumption) of O₂ and air-sea O₂ exchange are the only processes affecting O₂, i.e., mixing is ignored, then gross PP equals:

PPg = G*O₂^sat *[D¹⁷O –D¹⁷O_eq] / [D¹⁷O_photo – D¹⁷O] (1)

Where G is air-sea gas transfer rate, O₂^sat is the O₂ concentration expected in equilibrium with the atmosphere, D¹⁷O_eq is the D¹⁷O expected for dissolved O₂ in equilibrium with the atmosphere (+16 per meg) and D¹⁷O_photo is the D¹⁷O of photosynthetically produced O₂ (+249 per meg).

The rate of net primary productivity (PPn) is obtained by simultaneously measuring the O₂ saturation levels due to biological activity along with D¹⁷O. (In this case, PPn represents the gross primary production minus community respiration.) The biologically produced O₂ saturation is determined from measurements of O₂ and Ar saturation levels where the Ar saturation represents the proportion of the O₂ saturation that is due to physical processes (gas exchange, warming, mixing). The biologically produced O₂ saturation is essentially represented by the difference between the O₂ and Ar saturation levels (Emerson et al., 1991). If air-sea O₂ gas exchange is the only process affecting O₂ concentrations in the mixed layer, then the O₂ and Ar saturation levels would both be 100%. If gross primary production exceeds net community respiration, i.e., PPn>0, then the O₂ saturation exceeds Ar saturation and vice-versa. Bender (2000) shows graphically the relationship between D¹⁷O and the degree of O₂ saturation (expressed as O₂/O_2sat) as a function of the ratio of net to gross PP (PPn/PPg) for a mixed layer where mixing effects were negligible (Fig. 3).

(Fig. 3)

Our recently funded project focuses on making Δ¹⁷O measurements several times during an annual cycle at the ALOHA time series station near Hawaii. This station is the site of a decade long biogeochemical study of carbon cycling in the subtropical N. Pacific. In particular, monthly measurements of primary production rates (using ¹⁴C) and particulate organic carbon sinking rates (from sediment traps) occur at ALOHA. We intend to begin these measurements in spring 2002. We will also make in vitro measurements of gross primary production rates using ¹⁸O labeled water following the procedures of Bender (199?). These measurements should begin towards the end of 2001. Our goal is to determine the rates of gross primary production using in situ Δ¹⁷O and in vitro ¹⁸O techniques and compare these rates to primary production rates estimated from ¹⁴C bottle measurements.

References

Bender, M.L. 2000. Tracers from the sky. Science 288: 1977-1978.

Bender M., J. Orchardo, M-L. Dickson, R. Barber and S. Lindley. 1999. In vitro O₂ fluxes compared with ¹⁴C production and other rate terms during the JGOFS equatorial Pacific experiment. Deep-Sea Res. 46: 637-654.

Luz, B., E. Barkan, M.L. Bender, M.H. Theimens and A. Kristie. 1999. Triple isotope composition of atmospheric oxygen as a tracer of biosphere productivity. Nature 400: 547-550.

Luz, B. and E. Barkan. 2000. Assessment of oceanic productivity with the triple isotope composition of dissolved oxygen. Science 288: 2028-2031.