Andy Ridgwell1 and Ursula Edwards2

1 School of Geographical Sciences, University of Bristol, Bristol, UK; 2Occidental Oil and Gas Corporation, Houston, Texas, USA

6.1 Introduction

The sequestering (locking up) of carbon in geological formations and removal of carbon dioxide (CO2) from the atmosphere is not a unique, human-driven invention thought up for ameliorating (reducing) the degree of greenhouse gas-driven climate change in the future. CO2 has been spewing from volcanoes on land and the spreading ridges of the ocean throughout geological time. Yet, CO2 levels are not thought to have risen inexorably since the Earth was formed some 4.5 billion years ago. Quite the opposite - geological evidence suggests that we live in a period in which atmospheric CO2 concentrations are probably amongst the lowest to have occurred on Earth, at least for the past 600 million years (Royer et al., 2004) (Fig. 6.1a). Indeed, compared to the 'pre-Industrial' atmosphere (i.e. immediately prior to the industrial revolution c. 1765 and the onset of industrialization and increasingly rapid fossil fuel consumption), which was characterized by a CO2 concentration of 278 ppm (Enting et al., 1994), geological periods such as the Jurassic (200-145 million years ago) and Devonian (416-359 million years ago) saw about ten times as much carbon residing in the atmosphere. The ocean carbon reservoir would also have been much larger at times in the past (Ridgwell, 2005)

(Fig. 6.1b). But where has this carbon gone? There must exist greenhouse gas sinks that are able not only to sequester large amounts of carbon but also to keep a tight hold of it for extended periods of time. The amount of carbon stored in vegetation and soils has varied significantly through time, with as much as ~1000 Pg C less at the height of the last glacial period associated with major shifts in vegetation type and coverage, compared to the ~2200 Pg C in the modern terrestrial biosphere (Fig. 6.2). However, the fall in atmospheric CO2 between the Devonian and Carboniferous (369-299 million years ago) equates to ~5000 Pg C reduction in carbon stored in the atmosphere (and a considerably greater reduction in the ocean inventory) (Fig. 6.1a). It is not easy to envisage how the terrestrial biosphere could possibly have accommodated this increase in carbon storage. The terrestrial biosphere is also not a 'safe' long-term store for carbon. For instance, the devastation wrought by the Indonesian wildfires of 1997 has been estimated to have resulted in the release of 0.8-2.6 Pg C (Page et al., 2002). This is equivalent to 13-40% of the annual emissions from anthropogenic fossil fuel combustion, and could help explain why the growth rate of CO2 in the atmosphere approximately doubled during the 1997-1998 period (Schimel and Baker, 2002).

©CAB International 2007. Greenhouse Gas Sinks (eds D.S. Reay, C.N. Hewitt, K.A. Smith and J. Grace)

Phanerozoic

Pre-Cambrian

Cenozoic

Mesozoic

Paleozoic

Cenozoic

Mesozoic

Paleozoic

|P9

-1--1-r'

i—H

i

1 1

T-1--

Nitrogen Deposition

4000 3000 2000

Age (million years ago)

Fig. 6.1. Evolution of atmospheric and oceanic carbon reservoirs through time. (a) Phanerozoic evolution of atmospheric CO2 reconstructed from proxy records (Royer et al., 2004). The filled squares show the data binned into intervals of 20 million years, with one standard deviation of the error shown as a vertical black line for each point. The raw proxy data are plotted as open circles. There are two vertical scales; atmospheric concentration (right) and the corresponding total carbon inventory (left). (b) Model-estimated evolution of the ocean carbon reservoir (Ridgwell, 2005). The vertical scales are mean dissolved inorganic carbon (DIC) concentration (right) and total ocean carbon inventory (left). The horizontal lines in both panels indicate the size of the present-day carbon inventories. The geological timescale abbreviations for the periods are: D, Devonian; C, Carboniferous; Pr, Permian; J, Jurassic; N, Neogene; P, Paleogene; K, Cretaceous; S, Silurian; O, Ordovician; e, Cambrian; P-e, Pre-Cambrian.

Geological period 7000

4000 3000 2000

2000

0 600

Age (million years ago)

Fig. 6.1. Evolution of atmospheric and oceanic carbon reservoirs through time. (a) Phanerozoic evolution of atmospheric CO2 reconstructed from proxy records (Royer et al., 2004). The filled squares show the data binned into intervals of 20 million years, with one standard deviation of the error shown as a vertical black line for each point. The raw proxy data are plotted as open circles. There are two vertical scales; atmospheric concentration (right) and the corresponding total carbon inventory (left). (b) Model-estimated evolution of the ocean carbon reservoir (Ridgwell, 2005). The vertical scales are mean dissolved inorganic carbon (DIC) concentration (right) and total ocean carbon inventory (left). The horizontal lines in both panels indicate the size of the present-day carbon inventories. The geological timescale abbreviations for the periods are: D, Devonian; C, Carboniferous; Pr, Permian; J, Jurassic; N, Neogene; P, Paleogene; K, Cretaceous; S, Silurian; O, Ordovician; e, Cambrian; P-e, Pre-Cambrian.

Clues as to the ultimate fate of CO2 released to the atmosphere lie in the rocks around us. Fossil fuel deposits such as coal measures, oil and gas reservoirs, as well as oil shales and other organic matter rich sedimentary rocks, all hold substantial quantities of carbon (Fig. 6.2). The relationship of these reservoirs to atmospheric CO2 is conceptually fairly straightforward - sequestration of organic matter in geological formations must result in less carbon in the atmosphere and ocean. Past increases in organic carbon burial driven by evolutionary and tectonic factors have been linked to decreases in atmospheric CO2, particularly the CO2 'trough' during the Carboniferous and Permian periods (Berner, 1990) (Fig. 6.1a). We are all too familiar with the converse link: the burning of deposits of ancient carbon and increasing CO2 concentrations in the atmosphere. However, if the rate of burial of organic matter were to increase in the future, perhaps in response to climate change, there would presumably be an additional removal of fossil fuel CO2 from the atmosphere. Is this likely to occur, and how important might this be? To answer this question we are going to have to look at how organic matter is deposited and preserved in accumulating sediments.

The formation of carbonate rocks such as limestones (the remains of ancient reefs) and

Atmospheric CO2

Terrestrial biomass

Marine biomass ar>o

"Oil and gas reservoirs

Dissolved inorganic carbon

Marine biota

3Gt

C

Organic matter in surface sediments

150Gt

C

Terrestrial vegetation

600Gt

C

Oceanic dissolved organic matter

< 700Gt

C

Atmosphere

750Gt

C

Humus, litter, and peat

1,600Gt

C

Carbonate in surface sediments

2,500Gt

C

Methane clathrates

3-10,000Gt

C

Recoverable coal, oil and gas

5-10,000Gt

C

Dissolved inorganic carbon

39,000Gt

C

Dispersed carbon (e.g. oil shales)

10,000,000Gt

C

Carbonate rocks

40,000,000Gt

C

Fig. 6.2. Estimated inventories of the various carbon reservoirs on Earth. (Adapted from Kump et al., 2004.) Units are in Gt C; 1 Gt C = 1 Pg (1015g) of carbon or about 8.3° x 1013 mol carbon. To put this into some perspective, the 750 Gt C in the atmosphere is equivalent to an average concentration of CO2 of 351 ppm, and the average emission of anthropogenic CO2 (from fossil fuels and cement production) during the 1990s was 6.3 Gt C/year. (From Houghton et al., 2001.)

chalks (the microscopic shells of dead calcifying marine plankton) also represents a sink of carbon from the Earth's surface. This turns out to be an even more important reservoir of carbon than fossil fuels and other forms of ancient organic matter (Fig. 6.2). The relationship between carbonate deposition and atmospheric CO2 is also less straightforward - when marine carbonates are precipitated from solution, the concentration of CO2 in the atmosphere actually goes up. How the burial of carbonate can at the same time be a geologic sink for carbon needs some explanation. We also need to quantify how this sink might change in the future, and whether it will be important on the (century and shorter) timescales that matter to us most.

In this chapter we examine the role played by 'geologic' sinks for fossil fuel CO2 emitted to the atmosphere. We tell the inorganic (carbonate) carbon and organic carbon sides of the story separately (in Sections 6.2 and 6.3, respectively), and look in detail at the underlying mechanisms involved. We also consider how human (anthropogenic) activities and climate change may perturb these processes in ways that will affect the rate at which fossil fuel CO2 is removed from the atmosphere. Finally, we provide a summary and perspective in Section 6.4. Box 6.1 contains a basic '101' tutorial on carbonate chemistry and the key geochemical reactions involved in the geologic carbon sink (see Zeebe and Wolf-Gladrow, 2001 for a more detailed primer).

Box 6.1. Carbonate chemistry '101'.

The mineral calcium carbonate (CaCO3) has a crystal lattice consisting of calcium ions (Ca2+) ionically bound to carbonate ions (CO3-). The lattice can take one of several different 'polymorphic' forms (i.e. the same chemical composition but different crystalline structure) such as calcite, or a higher symmetry aragonite phase. Calcite is the more abundant of the two polymorphs that are biologically precipitated in the open ocean. Because it is also more thermodynamically stable than aragonite it is the phase responsible for almost all carbonate burial in the deep sea. In contrast, aragonite is abundant amongst shallow water carbonates (e.g. corals). Biogenic carbonates are not always pure CaCO3 and a range of substitutions of magnesium (Mg2+) for Ca2+ are possible in the crystal structure to give natural carbonates a generic formula: Mgx x Ca(1_x) x CO3. We will, however, focus on the more abundant CaCO3 end member here.

Precipitation of calcium carbonate may be described by the following reaction:

Because Ca2+ has a 'residence' time in the ocean counted in millions of years, we can assume that the concentration of Ca2+ does not change on the (100-100,000 year) timescales we are interested in, and that ocean mixing homogenizes its concentration throughout the ocean. The bicarbonate ion (HCO-) required in the precipitation reaction is formed through the hydration and dissolution of CO2 gas (CO2(g)) to form a proton (H+) and a bicarbonate ion (HCO-):

(The first, hydration step is the formation of carbonic acid (H2CO3), but this is only present in very small concentrations and is commonly ignored.) A fraction of the HCO- dissociates to form a carbonate ion (CO3-) and another proton:

The sum total of carbon in all its dissolved (inorganic) forms (i.e. CO2(aq) + HCO- + CO3-) is collectively termed dissolved inorganic carbon (DIC). The relationship between these different forms of DIC, and to hydrogen ion (H+) concentrations and to pH is shown in Fig. 6.3.

When CaCO3 is precipitated from solution, although the total sum of DIC is reduced, the remaining carbon is re-partitioned in favour of the CO2(aq) species. One way of thinking about this is in terms of removing CO3- and shifting the aqueous carbonate equilibrium reaction:

to the left to compensate. The counterintuitive consequence of this is that the precipitation of carbonate drives an increase in the partial pressure of CO2 (pCO2) in the surface ocean, despite there being a reduction in total carbon (DIC). (pCO2 is the variable that determines the exchange of CO2 between ocean and atmosphere - if atmospheric pCO2, which is equal to the CO2 molar ratio at a surface pressure of 1 atmosphere is greater than the pCO2 of the surface ocean, there will be a net transfer of CO2 from the atmosphere to the ocean, and vice versa.)

Whether CaCO3 precipitates or dissolves is dictated by the stability of its crystal structure relative to the ambient environmental conditions. This can be directly related to the concentrations of Ca2+ and CO3- and written in terms of the 'saturation state' (also known as the solubility ratio) W of the solution, defined as:

W = (Ca2+) x (CO32-) / Ksp where Ksp is a solubility constant. The precipitation of calcium carbonate from sea water is thermodynamically favourable when W is greater than unity. Conversely, CaCO3 will tend to dissolve at W < 1.0. In addition to the concentrations of Ca2+ and CO3-, depth is also important because Ksp scales with increasing pressure (as well as with decreasing temperature). Thus, the greater the depth in the ocean, the more the ambient environment will tend to be undersaturated (i.e. W < 1.0) and the less likely that carbonate will be present in the sediments. . , continued

Box 6.1. Continued

Box 6.1. Continued

Inorganic Species Seawater

Fig. 6.3. The concentrations of the dissolved carbonate species as a function of pH (referred to as the Bjerrum plot; cf. Zeebe and Wolf-Gladrow, 1999): dissolved carbon dioxide (CO2(aq)), bicarbonate (HCO-), carbonate ion (CO2-), hydrogen ion (H+) and hydroxyl ion (OH-). At modern sea water pH, most of the dissolved inorganic carbon is in the form of bicarbonate.

Fig. 6.3. The concentrations of the dissolved carbonate species as a function of pH (referred to as the Bjerrum plot; cf. Zeebe and Wolf-Gladrow, 1999): dissolved carbon dioxide (CO2(aq)), bicarbonate (HCO-), carbonate ion (CO2-), hydrogen ion (H+) and hydroxyl ion (OH-). At modern sea water pH, most of the dissolved inorganic carbon is in the form of bicarbonate.

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