Limits to Marine Life

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Science  17 Apr 2009:
Vol. 324, Issue 5925, pp. 347-348
DOI: 10.1126/science.1170756

Ocean chemistry is currently undergoing enormous change from the twinned impacts of higher carbon dioxide (CO2) concentrations from fossil-fuel burning (1), inducing ocean acidification (2, 3), and of rapidly declining mid-water oxygen (O2) concentrations. The decline in O2 results from lower sea-surface O2 concentrations, reduced ventilation of the mid-water from ocean warming (4, 5), and local eutrophication events, all of which lead to an expansion of oceanic dead zones. The reduced ventilation further elevates CO2 concentrations at depth, because the decline in O2 is accompanied by the equivalent respiratory CO2 (6); as a result, ocean acidification penetrates more rapidly to lower depths than it would due to the fossil-fuel signal alone. Can the effects of these changes on marine life be quantified on the basis of existing data, and if so, how does one quantify them?

Initial concerns over ocean acidification focused on reduced calcification in coral reefs and other calcareous organisms (7, 8), but other concerns soon arose. Elevated dissolved CO2 concentrations may impose a physiological strain on marine animals, impairing performance and requiring energy that would otherwise be used for locomotion, predation, reproduction, or coping with other environmental stresses such as rising temperatures. However, there is as yet no formal way to estimate this impact or to relate observed oceanic chemical change to the physiological limits for marine organisms.

Ocean scientists today define the limits to aerobic life in the sea in terms of a minimum dissolved O2 concentration (5), typically ∼5 μM, below which it is inefficient for aerobic microbes to consume dissolved O2; instead, the microbes turn to other electron acceptors such as IO3, Mn(IV), and NO3 (9). For higher animals, “dead zones” are defined as regions where normal respiration is greatly limited and the expenditure of effort is physiologically constrained, but there is no precise, universally accepted definition that would allow a common limit to be used when mapping changing conditions. In writing this limit in terms of dissolved O2 [or oxygen partial pressure (pO2)] alone, ocean scientists typically ignore the CO2 side of the respiration equation, on the unspoken assumption that pCO2 levels are low and are inversely proportional to the O2 concentration via bacterial oxidation of marine organic matter. However, this may no longer be the case as atmospheric CO2 concentrations rise and reset ocean chemical relations.

A simple way to approach this problem is to define the basic oxic respiration equation

Corg + O2 → CO2 (1)

and from this write out the free-energy relation

ΔG = ΔG° − RT · ln{[f CO2]/[Corg][f O2]} (2)

Here, ΔG° is the Gibbs free energy at standard conditions, R is the universal gas constant, T is temperature, and f is fugacity. From this equation, we can see that for all food sources there is a common term: the natural logarithm of the ratio of the gas fugacities. By substituting partial pressures for fugacities, log10 for the natural logarithm, and inverting the ratio to eliminate the minus sign, we obtain the expression log10 (pO2/pCO2), which provides a simple numerical constraint that is linearly related to available energy. We define this as the respiration index (RI), which may prove useful for estimating the physiological limits of deep-sea animals.

For a specific example, consider a station in the eastern tropical Pacific that is characteristic of the very large suboxic regions of the oceans. Here, the dissolved O2 concentrations at depths between 300 and 600 m decline almost to zero, and large-scale reduction of nitrate to nitrite occurs. In this region, the calculated RI ranges from just below zero to 2 or more, depending on the ratio of pO2 to pCO2 (see the figure). Field data (10) suggest that denitrification begins to occur at RI ≈ 0.4 to 0.7, and this ratio likely sets the limit for aerobic respiration of higher animals. Actual limits will be species dependent and remain to be determined (see the figure for some hypothetical limits).

Expanding dead zones.

An example of respiration stress at a station in the eastern tropical Pacific (WOCE P16C Sta 413: 13°01.75′N, 91°45.60′W). (A) The calculated pCO2 rises with increasing atmospheric CO2 concentration. The preindustrial profile was calculated from the modern data by removing the anthropogenic CO2 at constant alkalinity. The projections (two times preindustrial and three times preindustrial) were calculated by determining the stepwise change in total CO2 in the sea surface for each case and then propagating this change throughout the ocean. (B) Calculation of the respiration index with depth (RI) reveals the existence of a formal dead zone for aerobic life, where RI ≤ 0 (gray band). However, even at RI = 0.0 to 0.4 (red band), aerobic respiration is not observed. Bacteria appear to set the practical limit for all aerobic respiration at RI = 0.4 to 0.7 (orange band). Some marine animals can tolerate RI = 1 or slightly less, but others cannot (yellow band). With increasing atmospheric CO2 concentrations, dead zones for aerobic life will grow in size. Rising ocean temperatures will further exacerbate the growth of the dead zones by decreasing oxygen saturation.

What is of concern is the impact of rising oceanic CO2 concentrations on this ratio. Present-day pCO2 at 500 m depth at this site is about 1000 μatm, but an increase of +280 parts per million by volume CO2 to the atmosphere and surface ocean translates into a far greater change at depth. As surface sea water is transferred to depth, its buffer capacity is reduced by the acidic components of the normal Redfield cycle (6). A doubling of surface-water pCO2 leads to a doubling or more of pCO2 at depth (see the figure) due to the different geochemistry of the deeper water masses. From ocean equilibration with a doubled CO2 atmosphere, the pCO2 at the example station at 500 m depth will rise to 2500 μatm and possibly higher. Such levels have not been considered previously in many of the models designed to predict the status of the future ocean.

This simple example uses the fossil-fuel CO2 signal alone, thereby greatly understating the case. The calculation assumes constant temperature, but oceanic warming is taking place and will drive the in situ pCO2 higher. The calculation also assumes unchanging pO2 levels, but deep-water O2 concentrations are steadily declining. This affects the RI both through reducing pO2 and through the associated increase in respiratory CO2. Thus, we may anticipate a very large expansion of oceanic dead zones.

The expansion of dead zones also has other chemical side effects. The major redox cycles of the chemical elements are microbially driven, and thus an increase in production of N2O—also a greenhouse gas—at depth seems likely, although any release of this to the atmosphere would be greatly limited by oceanic processes of mixing and consumption. Other redox species may serve as important tracers of the processes at work as these changes occur, with the IO3 → I system being the most sensitive indicator.

For the vast areas of the ocean that are well-oxygenated, the rise in oceanic CO2 concentrations will exert a negligible effect on the normal aerobic functioning of adult marine animals. However, based on our redefinition of dead zones, it is clear that even if oxygen levels do not decline, the oceanic dead zones will still expand as a result of rising CO2 concentrations; with global warming reducing the oxygen levels as well, the combined effect will be severe.


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