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Ocean Storage 12 Ocean storage 12.1 Introduction As described in Chapter 2, the world’s oceans contain an estimated 39,000 Gt-C (143,000Gt-CO2), 50 times more than the atmospheric inventory, and are estimated to have taken up almost 38% (500Gt-CO2) of the 1300Gt of anthropogenic CO2 emissions over the past two centuries. Options that have been investigated to store carbon by increasing the oce­ anic inventory are described in this chapter, including biological (fertilization), chemical (reduction of ocean acidity, accelerated limestone weathering), and physical methods (CO2 dissolution, supercritical CO2 pools in the deep ocean). 12.2 Physical, chemical, and biological fundamentals 12.2.1 Physical properties of CO2 in seawater The behavior of CO2 released directly into seawater will depend primarily on the pressure (i.e., depth) and temperature of the water into which it is released. The key properties are: l The liquefaction pressure at a given temperature: the point at which with increasing pressure, gaseous CO2 will liquefy l The variation of CO2 liquid density with pressure, which determines its buoyancy relative to seawater l The depth and temperature at which CO2 hydrates will form Saturation pressure At temperatures of 0–10°C, CO2 will liquefy at pressures of 4–5MPa, corre­ sponding to water depths of ~400–500m, with a liquid density of 860kg per m3 3 at 10°C and 920kg per m at 0°C. At this depth liquid CO2 will therefore be positively buoyant, and free liquid droplets will rise and evaporate into gas bubbles as pressure drops below the saturation pressure. Buoyancy Figure 12.1 shows the densities of CO2 and seawater versus depth for a range of ocean conditions. Liquid CO2 is more compressible than water and becomes © 2010 Elsevier, Inc. All rights reserved. Doi:10.1016/B978-1-85617-636-1.00012-2. 268 Carbon Capture and Storage neutrally buoyant at depths of 2500–3000m. Below this depth range, released liquid droplets will sink to the ocean floor. Hydrate formation Hydrates are a form of weakly bound molecular complex in which a guest molecule is trapped within a cage of water molecules. The cage of water molecules takes the form of a 12- or 14-faced hedron (dodecahedron or tetrakaidecahe­ dron) having pentagonal or hexagonal faces with an oxygen atom at each ver­ tex, similar to normal ice, as illustrated in Figure 12.2. The stoichiometry of this complex is approximately (CO2: 6H2O), depending on the temperature and pressure of formation. Figure 12.1 CO2 and seawater density versus depth Figure 12.2 CO2 hydrate structure Ocean storage 269 CO2 hydrates can form in all seawater below 400m, as long as there is sufficient CO2 present, and as shallow as 150m in polar regions where water temperature is close to 0°C. A CO2 gas bubble or liquid drop at hydrate stable depths will rapidly form a crystalline hydrate skin, which will limit the rate of dissolution of CO2 into the surrounding seawater. Naturally occurring hydrates have been observed forming around CO2-rich gas bubbles venting from hydro­ thermal vents at 1500m depth in the mid-Okinawa Trough. 3 The density of solid CO2 hydrate is 1.11 · 10 kg per cubic meter, and it will therefore sink in seawater, which has a density of 1.03 · 103 kg per cubic meter. However, the density of a gas hydrate mass formed from gas bubbles or liquid droplets wrapped in a hydrate skin will depend on the gas or liquid density, as well as the skin thickness and droplet size. 12.2.2 The ocean carbon cycle As described in Chapter 1, the deep ocean is by far the largest sink in the global carbon inventory, and two interlinked cycles—one physical and one biological—drive the fluxes into and out of that sink. The solubility pump The solubility pump, or physical pump, is the term used to describe the physi­ cal and chemical process by which CO2 is transported from the atmosphere to deep ocean waters. The process is illustrated in Figure 12.3, together with the interlinked biological pump described below. The two key elements of the solubility pump are the dissolution of CO2 into surface waters and the successive downwelling of surface water and upwelling Figure 12.3 The ocean carbon cycle: solubility and biological CO2 pumps 270 Carbon Capture and Storage of deep water via the global thermohaline circulation system. The surface layer of the ocean is well mixed by wind and wave action (hence also known as the mixed layer) and CO2 in the atmosphere reaches thermodynamic equilibrium with dissolved CO2 in this layer on a time scale of weeks to months, depending on mixed layer depth and wind conditions. The equilibrium solubility of a gas is directly proportional to its partial pressure (Henry’s law): C P /K PK CO 2 CO 2 HCO 2 CO 2 CO 2 (12.1) where CCO2 is the dissolved CO2 concentration, PCO2 is the partial pressure of CO2 in the atmosphere, KHCO2 is the Henry’s law constant, and KCO2 is the solubility of CO2. The solubility KCO2 is a function of pressure, salinity, and acidity (pH) and, most importantly, of temperature, with solubility increasing from 0.03 mol-CO2 per kg at 20°C to 0.05 mol-CO2 per kg at 10°C (for typical current atmospheric and oceanic conditions). The dissolution of CO2 in water initially forms carbonic acid (H2CO3), which will dissociate into bicarbonate and then carbonate ions, the partition­ ing depending on the water temperature and alkalinity. The total concentration 2 of these species—dissolved free CO2 (CO2 (aq)), carbonate ions (CO3 ), and bicarbonate ions (HCO3 )—is known as the dissolved inorganic carbon (DIC). The interactions that establish this balance are as follows: 2 CO 2 H 2 O ↔ H 2CO 3 ↔ HCO3 H ↔ CO3 2H (12.2) At current average mixed layer conditions, DIC is 91% bicarbonate and 8% carbonate ions, with the remaining 1% being dissolved free CO2 plus carbonic acid. The DIC content of surface waters varies from a low of 1850mol-C per kg in the Bay of Bengal to a high of 2200mol-C per kg in the Weddell Sea. The second component of the solubility pump is the formation of deep waters through meridional overturning that occurs in the North and South Atlantic Oceans (Figure 12.4). A parcel of water moving northward in the Gulf Stream experiences strong evaporative cooling due to wind action, resulting in increasing uptake of CO2 from the atmosphere as a result of its temperature-dependent solubility, as well as an increase in salinity and den­ sity as a consequence of evaporation. By the time a parcel of water reaches the Norwegian Sea, the density becomes high enough to cause it to sink, forming North Atlantic Deep Water (NADW). This water flows back in a southerly direction into the abyssal Atlantic basin. A similar process occurs in the sub-Antarctic South Atlantic, forming Antarctic Bottom Water (AABW). The total rate of downwelling is estimated at 30–40 · 106 m3 per day, and it is roughly evenly distributed between the North and South Atlantic locations. Ocean storage 271 Figure 12.4 Global thermohaline circulation Figure 12.5 Temperature versus depth in a tropical ocean Other than at these two downwelling locations, a stable layering or strati­ fication occurs in the ocean as a result of temperature and density gradients. This stratification can be seen in the tropical ocean temperature profile shown in Figure 12.5. A warmer, less dense upper layer is separated from the colder, denser deep water by an interval known as the thermocline, where temperature drops rapidly with depth. The thermocline thus marks the limit of the mixed layer, within which fairly rapid thermodynamic equilibrium is reached between the surface water and the atmosphere. Deep ocean waters, with temperatures down to 2°C and pressures 10MPa, are highly undersaturated, and it is this excess solubility that enables the ocean to provide a sink for increased atmospheric CO2 over a centennial time scale. 272 Carbon Capture and Storage The deep waters from the Atlantic Basin circulate westward around the Antarctic Ocean basin, branching off and subsequently upwelling in the Indian and Pacific Oceans as well as in the Southern Ocean. The time scale of this circulation system is millennial. Early estimates of the time required for down- welling surface water to fully displace the deep ocean volume based on 14C radioactive tracer measurements put this “ventilation time” of the deep ocean at 1400 years. Subsequent improved understanding of the processes involved has reduced this time scale to 250 years for the Atlantic and 550 years for the Pacific. When deep waters upwell, CO2 outgases to the atmosphere at the seawater– air interface since the partial pressure of CO2 in the seawater is higher than the partial pressure in the atmosphere. In preindustrial times this cycle had reached an equilibrium, with average surface and deep waters containing 2000 and 2100mol-C/kg of DIC respectively. The anthropogenic increase in CO2 partial pressure in the atmosphere, from 280ppm in preindustrial times to the current 380ppm, has resulted in an increase in the average DIC of surface water to 2080mol-C per kg. As a result, the ocean has already sequestered an estimated 500Gt of anthropogenic CO2 emissions over the past two centuries. However, the ability of the surface water to take up additional CO2 does not increase linearly as atmospheric CO2 concentration increases. This is because, as dissolved CO2 and therefore carbonic acid increases, the preferred reaction is: 2 CO 2(aq) H2 O CO3 → 2HCO3 (12.3) This reaction consumes carbonate ions in a process known as carbonate buffering, and reduces the impact on ocean acidity (pH) that would otherwise arise, since it converts carbonic acid to alkaline bicarbonate ions.
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