Peter G Brewer

It is now 25 years since Marchetti (1977) first suggested bypassing atmospheric disposal of some fraction of industrial CO2 emissions and using direct deep-ocean disposal as one means of ameliorating climate change. After all, the alkalinity of the ocean already provides the dominant long-term sink for atmospheric CO2, and deep-ocean injection may logically be seen as simply accelerating a "natural" process. Behind this apparently simple suggestion lies great complexity, fascinating science, and strongly held opinions. The scale of modern CO2 fluxes, and the interest in finding safe and economically viable ways to avoid "dangerous anthropogenic interference with climate," has led to new efforts to investigate this possible solution. Until very recently, only cartoon sketches were available to describe the possibilities of this field (Hanisch 1998). Now, important and challenging small-scale field experiments (Brewer et al. 1999) and new ocean modeling studies (Drange et al. 2001; Caldeira et al. 2002) have been conducted, and they illuminate both the possibilities and the challenges facing this strategy. A significant international scientific literature now exists (Handa and Ohsumi 1995), and the field has become established as a growing part of ocean science.

There are many variants of proposed ocean CO2 disposal strategies. For example, Caldeira and Rau (2000) proposed using limestone to neutralize saturated solutions of captured CO2, followed by shallow ocean disposal of the resulting solution. All of these concepts will have to be subjected to experimental test, and much of this work remains to be done. Early experiments on small-scale CO2 injections form the basis of this background chapter.

The scale of surface ocean CO2 uptake from the atmosphere is now so large that any distinction between a "natural" process and an industrial policy of only slightly indirect surface ocean disposal is hard to maintain. The average rate of surface ocean CO2 uptake for the 1980s and 1990s was approximately 2.0 petagrams of carbon per year (PgC y-1). This is 21 million tons of CO2 per day, and fluxes of such scale will have bio-geochemical impacts. Surface ocean waters are already >0.1 pH units lower than in preindustrial times (Brewer 1997), and if the IS92a "business-as-usual" scenario of the Intergovernmental Panel on Climate Change (IPCC) is followed, by the end of this century surface ocean carbonate ion concentrations will drop by 55 percent, with anticipated major effects on coral reef systems (Langdon et al. 2000) and calcareous plankton (Riebesell et al. 2000).

With this as background, it is clear that the enormous engineering demands would limit any industrial effort to dispose of CO2 by direct injection in the deep sea to a very small fraction of the massive current and future surface fluxes. Nonetheless, such techniques could be very useful as part of a portfolio of atmospheric stabilization strategies, if they are shown to be safe and cost effective. The criticisms leveled at this approach are essentially that it would be harmful to marine life (Seibel and Walsh 2001) and that it would inevitably increase the already large oceanic burden of fossil-fuel CO2.

Such comments at once expose scientific unknowns, for there are simply no standards to refer to here, and standards are urgently needed. How sensitive are marine animals to pH changes? If we permit atmospheric CO2 levels to rise to ~600 parts per million (ppm) and remain there, then surface ocean waters, and eventually all ocean waters, will experience an ~0.3 pH change from preindustrial levels. Is that an acceptable limit, and why? What burden of fossil fuel CO2 is "acceptable" for climate or geo-chemical or biological reasons? How might we view the trade-off between the direct effects of CO2 and the induced effects of climate and temperature? The focus of the ocean science community has been on observing the evolving fossil-fuel CO2 tracer signal, and these questions have simply (and astonishingly) not yet been posed. It is clear that the atmosphere, and therefore also the ocean, has experienced very large changes in CO2 over Phanerozoic time (Berner 1990). And in today's ocean there are large pH gradients between the Atlantic and Indo-Pacific oceans and within ocean basins. Many marine animals migrate vertically through large pH gradients each day. How these observations relate to the modern fossil-fuel signal remains to be tested.

Given these questions, the current series of deep-ocean CO2 injection experiments can be viewed not simply as disposal strategy tests, but also as controlled enrichment experiments that may allow us to mimic the elevated CO2 levels of a future ocean in much the same way that ecosystems are manipulated on land (DeLucia et al. 1999; Shaw et al. 2002).

The cost of CO2 capture dominates the economics of both geologic and oceanic disposal schemes. An excellent introduction to this problem is provided by the International Energy Agency's Greenhouse Gas Program ( The most widely used technology is scrubbing the gas stream with an amine solvent and regenerating the pure CO2 by heating the amine. This technique is routinely used in the food and beverage industries and for capturing CO2 for geologic disposal in the Sleipner Field off Norway. The costs of the capture process presently exert an ~25 percent energy penalty, and there are now strenuous efforts to improve this.

01 3

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""\n Situ

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" < 905m

Temperature (°C)

Figure 27.1. The phase behavior of CO2 in seawater, showing the gas-liquid and hydrate phase boundaries with a typical in situ P-T profile (from Brewer et al. 1999)

The Near-Field Fate of CO2 in Seawater

Ocean scientists are well versed in the thermodynamics and ocean distributions of the extraordinarily dilute (~2.2 millimolar oceanic CO2 system (Wallace 2001). But such concepts must be radically extended to investigate the science of deep-ocean CO2 injection. Figure 27.1 shows the phase diagram for pure CO2 superimposed on an oceanic pressure-temperature (P-T) scale (Brewer et al. 1999), with a temperature profile from a station off northern California overlaid.

The shaded area in Figure 27.1 indicates the zone in which CO2 will react with seawater to from a solid hydrate (CO2.6H2O), with profound changes in physical behavior. For the P-T profile indicated, CO2 gas will first form a hydrate at a depth of about 350 meters (m). The transition from gas to liquid occurs at about 400 m deep. For warm water regions (such as the Sargasso and Mediterranean Seas), these profiles will be shifted, but Figure 27.1 is generally applicable over much of the world's oceans. Liquid CO2 is highly compressible, and at depths < 2,750 m, it is less dense than seawa-ter; thus CO2 released will form a buoyant plume that will dissolve in the surrounding ocean (Alendal and Drange 2001). The dissolution rate of CO2 droplets is close to 3

Figure 27.2. A "frost heave" of CO2 hydrate on the sea floor at 3,600 m depth resulting from massive hydrate formation.

micromoles per square centimeter per second (|lmol cm-2 sec-1) (Brewer et al. 2002a), with the result that, for a release of small droplets, 90 percent of the plume is dissolved within 30 minutes and within about 200 m above the release point.

Below a depth of about 3,000 m, the high compressibility of liquid CO2 results in formation of a fluid of greater density than seawater, and a gravitationally stable release is possible. This finding has led to suggestions of storing liquid CO2 as a "lake" on the deep ocean floor. Here considerable complexity exists. The solubility of CO2 in seawater is so high (~ 0.8 molar at 1(C and 30 megapascals [MPa]) and the partial molal volume so low (~ 31 milliliters per mole [ml mol-1]) that a dense boundary layer can form (Aya et al. 1997), inhibiting mixing with the ocean above. Such a system would have much in common with naturally occurring pools of dense brines on the seafloor. The formation of hydrate in such a system, however, can occur spontaneously, with complex self-generated, fluid dynamic instabilities and resulting large volume changes from the incorporation of six molecules of water for every molecule of CO2 (Brewer et al. 1999). A spectacular example of this is shown in Figure 27.2, where an experimental pool of CO2 placed on the seafloor at 3,600 m depth has penetrated the upper sediments and formed a massive hydrate "frost heave."

It was earlier thought that the formation of a hydrate would possibly result in "per manent" storage of CO2 on the seafloor as a hydrate because the P-T conditions are so strongly favorable. The essential condition for hydrate stability, however, is equality of the chemical potential in all phases. Because seawater is so strongly undersaturated with respect to CO2, the hydrate will dissolve. The oceanic dissolution rates of both hydrate-coated CO2 droplets (Brewer et al. 2002b) and of the solid hydrates (Rehder et al. 2002) have been directly measured.

Possible technologies for future applications may include either pipeline or tanker injection techniques (Aya et al. 2003). With either approach, injection depth, physical state, or local pH perturbation can be optimized.

The Far-Field Fate of CO2 in Seawater

Once fossil-fuel CO2 injected into deep-ocean water becomes dissolved in the background ocean, it yields a tracer signal that can be modeled. Aumont et al. (2001) compared seven ocean models to explore the long-term fate of purposefully injected CO2. Essentially all show a strong correlation between depth of injection and efficiency of retention. For injection at 3,000 m ocean depth, overall retention efficiencies are ~ 85 percent over a 1,000-year timescale. Some surface ocean reexposure and reabsorption from the atmosphere is included in this estimate. For a 3,000 m injection, this fraction is about 25 percent of the total.

Experimental Activities

For many years, only model concepts and sketches were available to describe this field. Plans for a medium-scale experiment (several tons CO2) off the coast of Hawaii, and later off the coast of Norway, were frustrated by environmental permitting issues. This situation changed in 1996, with the successful adaptation of remotely operated vehicle (ROV) technology to contain and release small, controlled quantities of CO2 in the deep ocean for scientific study (Brewer et al. 1998). Rapid advances in technique then led to the ability to conduct experiments at great depth (Brewer et al. 1999). These experiments have now been extended to elegant biological response studies (Barry et al. 2002) and sophisticated chemical measurements (Brewer et al. 2002b).

The most recent experimental system used is a 56-liter (l), carbon-fiber, composite piston-accumulator (Hydratech, Fresno, CA), with a 23-centimeter (cm) outside diameter and 194-cm length, rated at 3,000 pounds per square inch gauge (psig), for ROV operation (Figure 27.3). Two tandem cylinder pumps, with capacities of 128 milliliters (ml) and 970 ml, provide power for accurate delivery of the contained CO2. The CO2 cools and compresses during descent to ocean depth, so that under typical conditions (900 psig on deck at 16°C; 1.6°C at 3,600 m), 45 l are available per dive for experimental purposes.

The availability of these experimental quantities permits tests of ecosystem responses

Figure 27.3. A 56-l carbon-fiber-wound CO2 delivery system installed on ROV Tiburon, showing end cap with gauges, delivery pumps on top, and valves to the left. The dispensing valve is attached to the robotic arm in front of the vehicle and is not shown (from Brewer et al. 2003).

to elevated deep-ocean CO2 levels. One recent experimental arrangement is shown in Figure 27.4. Here, a set of small CO2 "corrals" has been placed on the seafloor at a depth of 3,600 m, where the liquid CO2 is sufficiently dense to be gravitationally stable. Around these CO2 sources are arranged animal cages, and cores are taken to investigate the benthic infauna. Current meters, a time-lapse camera, and recording pH sensors complete the observing system.

Such experiments are at a very early stage, and rapid progress is expected. The challenges of carrying out field experiments in the deep sea should, however, not be underestimated. For example, the well-known tidal flows force the low pH plume from the CO2 corrals, so that the experimental cages are exposed to varying pH signals (Figure 27.5). It should be possible to advance experimental design to improve this.


Deep-ocean CO2 sequestration is clearly technically possible, although there are many variants in the details, and few of these have yet been explored thoroughly. The primary costs of any CO2 capture and disposal technology lie in the capture step, and this situation is the same for either geologic or oceanic disposal. Early ideas of "permanent"

Figure 27.4. A sketch of a deep-sea CO2 enrichment experimental site designed to investigate the response of marine organisms to locally elevated CO2 levels. Note that deep ocean CO2 levels will rise even in the absence of any active carbon sequestration program (from Barry et al. 2002).

Figure 27.5. The pH signals recorded at distances of 1 m, 5 m, and 50 m from a CO2 source placed on the seafloor. The effects of electrode drift have been removed, and the data sets are offset for visual clarity. The background oceanic pH is about 7.6. The effect of the tidal velocity ellipse is to advect the CO2 plume past each sensor with about a 12.4-hour period (from Barry et al. 2002).

disposal as a hydrate on the seafloor are not realistic because the hydrate will readily dissolve in the unsaturated ocean water. Concepts of a "lake" of CO2 on the seafloor, with a stable, dense boundary layer above, remain to be tested, but these would be locally specific solutions where seafloor topography permits. In most locations, CO2 would quickly become dissolved in seawater and transported as a tracer plume by the abyssal circulation. The mean ventilation time of the Atlantic Ocean is ~250 years, and that of the Pacific Ocean ~550 years (Stuiver et al. 1983). Models show the expected Antarctic atmospheric reexposure of CO2 on approximately these timescales. Much of the CO2 is reabsorbed, with the result that, over a timescale of about 1,000 years, the injection is some 85 percent efficient.

The primary concerns over oceanic injection are the possible effects on marine life, and the fact that it will add to the already substantial ocean fossil-fuel CO2 burden. There are as yet no standards to refer to here, and these are urgently needed. The current invasion rate of fossil-fuel CO2 from the atmosphere to the surface ocean now approximates 20 million tons per day. This flux has already changed surface ocean pH, and a 0.3 pH reduction in ocean waters can be expected if atmospheric CO2 is held at about 600 ppm. Thus, deep-sea ecosystems will inevitably experience change, and it is likely that any disposal strategy will be of significantly smaller scale than the surface invasion signal.

Release of CO2 directly into the ocean may be less harmful than release of CO2 into the atmosphere. In this case, if injected CO2 results in diminished atmospheric peak concentrations, then direct injection of CO2 in the ocean could reduce the overall adverse consequences of fossil-fuel burning. If oceanic injection simply adds to atmospheric releases, however, overall adverse consequences could increase. Thus, the ability of oceanic injection to contribute to diminished adverse consequences of fossil-fuel use depends both on the science and technology of oceanic CO2 injection and on the roles that that science and technology might play in our energy economy.


I acknowledge the David and Lucile Packard Foundation, the U.S. Department of Energy Ocean Carbon Sequestration Program, and an international research grant from the New Energy and Industrial Technology Organization for support of this research.

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