Balancing the Oceanic Calcium Carbonate Cycle: Consequences of Variable Water Column W

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1 Aquat Geochem (2011) 17: DOI /s ORIGINAL PAPER Balancing the Oceanic Calcium Carbonate Cycle: Consequences of Variable Water Column W Stephen V. Smith Jean-Pierre Gattuso Received: 6 May 2010 / Accepted: 19 August 2010 / Published online: 3 September 2010 Ó Springer Science+Business Media B.V Abstract The paired chemical reactions, Ca 2?? 2HCO 3 - $ CaCO 3? CO 2? H 2 O, overestimate the ratio of CO 2 flux to CaCO 3 flux during the precipitation or dissolution of CaCO 3 in seawater. This ratio, which has been termed w, is about 0.6 in surface seawater at 25 C and at equilibrium with contemporary atmospheric CO 2 and increases towards 1.0 as seawater cools and pco 2 increases. These conclusions are based on field observations, laboratory experiments, and equilibrium calculations for the seawater carbonate system. Yet global geochemical modeling indicates that small departures of W from 1.0 would cause dramatic, rapid, and unrealistic change in atmospheric CO 2. W can be meaningfully calculated for a water sample whether or not it is in equilibrium with the atmosphere. The analysis presented here demonstrates that the atmospheric CO 2 balance can be maintained constant with respect to seawater CaCO 3 reactions if one considers the difference between CaCO 3 precipitation and burial and differing values for w (both \1.0) in regions of precipitation and dissolution within the ocean. Keywords Calcium carbonate reactions Seawater carbonate buffering Geochemical cycles CO 2 exchange Electronic supplementary material The online version of this article (doi: /s ) contains supplementary material, which is available to authorized users. S. V. Smith (&) Centro de Investigación Científica y de Educación Superior de Ensenada (CICESE), Ensenada, Baja California, México svsmith@cicese.mx J.-P. Gattuso INSU-CNRS, Laboratoire d Océanographie de Villefranche, B.P. 28, Villefranche-sur-mer, Cedex, France J.-P. Gattuso Université Pierre et Marie Curie, Observatoire Océanologique, Villefranche-sur-mer, France

2 328 Aquat Geochem (2011) 17: Introduction It is well accepted by geochemists that CaCO 3 precipitation releases CO 2 to the atmosphere, while dissolution takes up atmospheric CO 2. These processes are commonly represented by the following equation Ca 2þ þ 2HCO 3 $ CaCO 3 #þco 2 "þh 2 O ð1þ attributed to Urey (1952), although Berner and Maasch (1996) observed that the equation has historical antecedents going back to the mid nineteenth century (Ebelmen 1845). In their classical study of CaCO 3 precipitation on the Bahama Banks, Broecker and Takahashi (1966) noted that each mole of CaCO 3 deposition is accompanied by a loss of about 0.6 mole of CO 2. Those authors concluded that the present CO 2 budget cannot be balanced. Although not explicitly stated, it is evident that the balance to which they referred is the 1:1 molar balance between CO 2 evasion and CaCO 3 precipitation from seawater in equilibrium with atmospheric CO 2, as implied by Eq. 1. It has become apparent that the imbalance noted by Broecker and Takahashi (1966) for the Bahama Banks is a general characteristic of surface seawater. Wollast et al. (1980) provided experimental CaCO 3 precipitation data from which the evasion/precipitation ratio can be calculated to be Those authors noted that the direction of CO 2 flux resulting from CaCO 3 precipitation was consistent with Eq. 1 (i.e., gas evasion from the water), but made no comment about the quantitative discrepancy between their experimental results and that equation. Smith (1985) concluded that this imbalance is an inherent characteristic of seawater (but not freshwater) and that it is attributable to the buffer capacity of seawater. Smith suggested that this tendency could be related to a coupling between organic carbon (OC) and inorganic carbonate (IC) reactions. Ware et al. (1991) referred to the evasion/precipitation ratio as the 0.6 rule. Smith and Veeh (1989) noted that, if there was not a coupling between IC and OC reactions in isolated, well-mixed, shallow marine ecosystems, CaCO 3 precipitation alone would drastically alter seawater ph, pco 2, and carbonate saturation state. If the IC and OC reactions work in concert, these variables are not dramatically perturbed. Those authors suggested that the coupled reactions with no net gas flux be represented as: Ca 2þ þ 2HCO 3 $ CaCO 3 #þch 2 O þ O 2 ð2þ All of the above-mentioned observations were made for seawater (salinity *35) at a temperature near 25 C and in approximate equilibrium with contemporary atmospheric pco 2 (*350 latm over the time span of the observations). Frankignoulle et al. (1994) confirmed and further generalized the above-mentioned observations, by deriving the evasion/precipitation ratio as a function of the equilibrium reactions for the aqueous CO 2 system. Those authors noted that lowered temperature and elevated pco 2 cause the ratio to climb towards 1.0. They predicted that an increase in the value of that ratio will be a consequence of rising atmospheric CO 2. They referred to the ratio as w, a notation which we will follow throughout the remainder of this paper. Other authors have noted this phenomenon (e.g., Sundquist 1993). Lerman and Mackenzie (2005) made it explicitly clear that they have incorporated W into the surface layer of their global ocean models for the carbon cycle. Most of the discussion of w has dealt, either explicitly or implicitly, with CaCO 3 precipitation from surface seawater in near equilibrium with some assumed atmospheric

3 Aquat Geochem (2011) 17: pco 2. In fact, until the derivation by Frankignoulle et al. (1994), calculation of W was based on the change of dissolved inorganic carbon (DIC) concentration in seawater at equilibrium with a specified pco 2 (typically, atmospheric) before and after total alkalinity (TA) reduction from a fixed amount of CaCO 3 precipitation (e.g., Smith 1985; Sundquist 1993). Indeed, that approach, rather than the Frankignoulle equations, was used for the present analysis. The direct calculation is done using buffer factors that quantify chemical changes in response to changes in DIC and TA (Frankignoulle 1994; Egleston et al. 2010). An important point to note is that W is an intensive property of seawater and applies equally well to seawater below the mixed layer as to water exchanging CO 2 with the atmosphere. For example, CaCO 3 dissolution below the mixed layer would cause the transfer of free CO 2 in the water back to HCO 3 in a proportion determined by W. The fact that W can be applied to water out of immediate contact with the atmosphere has interesting consequences which we will demonstrate. A comment by Berner et al. (1983) makes it apparent that it is important to clarify the function of W in the carbon cycle. Those authors stated a 10% drop in the addition of CO 2 to the atmosphere via oceanic CaCO 3 precipitation (all other fluxes remaining constant) would result in the complete removal of atmospheric CO 2 in only 30,000 years. Such rapid changes do not occur, and, thus, fluxes must be much more closely in balance. Berner et al. then asserted that there must be a balance in the CaCO 3 reactions, without reconciling the assertion with respect to the 0.6 rule. Our analyses below provide that reconciliation. Following the notation of Frankignoulle et al. (1994), it is convenient to re-write Eqs. 1 and 2 as follows: Ca 2þ þ 2HCO 3 $ CaCO 3 #þwðco 2 "þh 2 OÞþð1 wþðch 2 O þ O 2 Þ ð3þ The above equation seems to be a reasonable approximation of how W might work in a well-mixed, atmospherically equilibrated 1-box system such as the Bahama Banks (Broecker and Takahashi 1966) or Spencer Gulf (Smith and Veeh 1989). Under these circumstances, CaCO 3 precipitation, OC production, and CO 2 flux would be in balance. There is, however, a problem when applying Eq. 3 with a constant value for W to the vertically stratified ocean. 2 Calculations All calculations of CO 2 -related variables were carried out as spelled out by Lewis and Wallace (1998). We used the Excel spreadsheet developed by Pierrot et al. (2006), including the default options for the dissociation constants and pressure coefficients. It can be demonstrated that w (the variable of primary interest) is not particularly sensitive to the constants used, because w is calculated from differences in DIC and TA rather than absolute values. The only variable not calculated directly in original application of the Excel spreadsheet is w. For the analysis presented here, w was calculated by using the spreadsheet to simulate the DIC change which would occur if some amount of CaCO 3 were precipitated with a known starting pco 2, and the sample were then allowed to return to that starting pco 2. DIC, TA, salinity, temperature, pressure, phosphate, and silicate data were used in the spreadsheet. The simulation was performed as follows for each data profile.

4 330 Aquat Geochem (2011) 17: First, pco 2 and both calcite and aragonite saturation states (X) were calculated directly with DIC and TA as the master CO 2 variables. DIC and TA derived from these calculations were normalized to a salinity of 35 in the presentation of Fig. 2 (below), while the other variables are entered directly from that first set of calculations. Second, observed TA (not the normalized TA) was decreased by an arbitrary amount (200 leq kg -1, equivalent to a DIC change from CaCO 3 reaction of 100 lmol kg -1 ; any value could have been used). DIC was then calculated using this new TA and the previously calculated pco 2 as the master variables. Third, w was calculated as (DDIC gas )/DDIC ta (i.e., the DIC evasion to precipitation ratio), where DIC gas = DIC tot - DIC ta. The subscripts tot, gas, and ta denote the total DIC change, the gas equilibration change, and the alkalinity-associated DIC change, respectively, between the first and second set of calculations. The values of w generated agree with those calculated using the direct calculation method implemented in seacarb (Lavigne and Gattuso 2010). 3 Examples from the Central Pacific Ocean Profiles of CO 2 -related variables in the Central Pacific Ocean provide insight into the significance of w. Figure 1 shows the location of Station Aloha, an oceanographic station 100 km north of Oahu, Hawaii (*4,700 m water depth; hot/hot-dogs/interface.html; last accessed 7 August 2010). That figure also shows locations of profiles along a north south transect (*150 W) that we have constructed from data in Fig. 1 Index map showing location of Station Aloha (green circle) and the World Ocean Database transect (red circles). Data from these stations are presented in Figs. 2 and 3

5 Aquat Geochem (2011) 17: Fig average data for temperature, salinity, sigma-t, dissolved inorganic carbon, and TA at Station Aloha (22.75 N, W) and pco 2, from which we calculate X for calcite and aragonite, as well as w. DIC * and TA * represent the DIC and TA values normalized to S = 35 to remove conservative effects. O 2 data are included for comparison with DIC the World Ocean Database (WOD; builder.pl; last accessed 7 August 2010). These two data sets are discussed below. Figure 2 presents 20-year average conditions of water composition at Station Aloha. Nutrient data are not included in the figure, even though phosphate and silicate data were used in the calculations as they contribute to TA. O 2 data are included for visual comparison with the DIC data. All data for that station were averaged as follows. A log 10 scale was used for depth, and raw data (in m) were bin-averaged into intervals of 0.1 log 10 units in order to highlight both the compositional difference between the upper and lower water column and the composition gradients between the two. For these data, TA and DIC (measured to high precision at the station) were used as the master variables to partition the CO 2 system. Calculations include hydrostatic pressure effects on the CO 2 system. Also included are calculations of w and carbonate mineral saturation state (X ¼½Ca 2þ Š½CO 2 3 Š= K 0 sp ; where [] denotes concentration, and K0 sp is the mineral solubility product constant) for the two prominent carbonate minerals precipitated by planktonic organisms, low-mg calcite and aragonite. Salinity varies vertically. In order to visualize the relatively small, but geochemically important, non-conservative behaviour of DIC and TA relative to salinity, both DIC and TA were normalized to a constant salinity of 35 (DIC * and TA *, respectively). To clarify, DIC and TA were used for calculations presented, but DIC * and TA * are used in the figure for ease of visualization. With these normalizations, it is seen that DIC * is constant above *100 m (the longterm averages of the base of the euphotic zone and the mixed layer depth approximately coincide at this station) then increases between that depth and about 900 m. DIC * decreases slightly below 2,000 m. TA * is constant to a depth of approximately 300 m (well below the mixed layer and euphotic zone), increases to a depth of *2,000 m, and then is constant. Both net organic carbon production and net CaCO 3 precipitation obviously occur in the upper water column. Therefore, constant, low values for DIC * and TA * in the upper water column relative to the deep ocean represent the combined effects of uptake of both DIC

6 332 Aquat Geochem (2011) 17: and TA into CaCO 3, uptake of DIC into organic matter, CO 2 gas exchange with the atmosphere mediated by both organic and inorganic carbon reactions, and mixing of this water to constant composition. The surface water pco 2 averages 340 latm, about 20 latm below average atmospheric pco 2 over the duration of the sampling period. For purposes of this analysis, we may consider the surface ocean pco 2 to be in approximate equilibrium with the atmosphere at this station. Note that a 20 latm difference between air and surface water pco 2 would result in a change in W much lower than 0.01 (Frankignoulle et al. 1994). Net CaCO 3 precipitation (represented by constant and low TA * relative to the deep ocean) persists downward to about 300 m, while both respiration and calcification affect DIC *, pco 2, X, and w. Rising DIC * and TA * between 300 and 900 m represents the combined effect of decomposing both organic matter and CaCO 3, with no gas exchange with the atmosphere. Organic decomposition is complete by *900 m, and there is little further downward change in X. The dramatic pco 2 decrease below 900 m, a slight continuing rise in TA, and a slight decrease in w may represent some continuing CaCO 3 dissolution in the absence of significant organic oxidation. More importantly, compositional variations at these depths are likely to be affected by horizontal transport of deep water masses in the absence of further significant vertical fluxes and reactions. This advective effect is particularly evident in the rise of O 2 at depths greater than 900 m, representing deepwater lateral advection of O 2 -rich water. w and X are constant in the euphotic zone/mixed layer. w rises from 0.6 to 0.9 between 100 and 800 m, in response to both elevated pco 2 from oxidation of organic matter and diminished temperature. Below 900 m, w decreases slightly as pco 2 is lowered by CaCO 3 dissolution in the absence of additional organic oxidation products. It is clear that CaCO 3 precipitation primarily occurs under conditions of one value for W (*0.6, m depth), while dissolution occurs at another (*0.9, below 300 m). By this interpretation, most CaCO 3 dissolution occurs above 1,000 m. CaCO 3 dissolution in the water column above the calcium carbonate saturation depth (often called the lysocline; see Boudreau et al. 2010) is contrary to long-held beliefs dating back to the Challenger Expedition. However, the trend we observe and our interpretation of it are consistent with the growing body of recent evidence (e.g., Milliman and Droxler 1996; Milliman et al. 1999; Troy et al. 1997; Feely et al. 2002, 2004) that most CaCO 3 dissolution occurs in water shallower than 1,000 m. The calcium carbonate saturation depth apparently marks the depth above which, rather than below which, most water-column dissolution of CaCO 3 occurs. A transect constructed from WOD data along *150 W through the Pacific Ocean (37 profiles; data collected since 1991; TA and DIC as the master variables for CO 2 calculations) allows us to assess how typical Station Aloha is of the distribution of w throughout the Central Pacific Ocean (Fig. 3). The figure includes temperature ( C) and pco 2 (latm) as the major correlates for w (Frankignoulle et al. 1994). As anticipated, w is well predicted by temperature and pco 2 based on a multiple linear regression of W against these two variables for the entire set of WOD data used here: (w = *T? *pCO 2 ; R 2 = 0.97, n = 1,046). In broad terms, the trend of surface w values near 0.6 and deep values near 0.9 is borne out along the transect between about 30 N and 30 S. At higher latitudes, surface w rises to above 0.7, primarily because of diminished temperature. The North Pacific Gyre has slightly higher values of w near a depth of 1,000 m than the South Pacific Gyre (slightly

7 Aquat Geochem (2011) 17: Fig. 3 Transect constructed from 37 hydrographic stations near 150 W in the Central Pacific Ocean. By convention, the graph shows the latitude axis with north being positive. Data processed the same as for Station Aloha. The data were then gridded to 5 degrees latitude and 0.2 log units of depth to construct contours. The approximate position of Station Aloha is indicated. The three panels represent temperature (C), pco 2 (matm), and y above 0.9 in the North Pacific, slightly below 0.9 in the South). This difference reflects higher pco 2 at these depths in the North Pacific Gyre relative to the South Pacific. 4 Calcium Carbonate Balance in a 1-D, 2-Box Ocean For this analysis, we rely primarily on the CaCO 3 budget of Milliman and Droxler (1996). Those authors presented a thoughtful and detailed estimate of CaCO 3 production by both planktonic and benthic organisms. The analysis presented by those authors is consistent with estimates presented by Wollast and Mackenzie (1983), Lee (2001), and Feely et al.

8 334 Aquat Geochem (2011) 17: (2004), although differing in detail. In a simple two-box model, we put all of the CaCO 3 production into the surface ocean box (above 100 m depth) and all of the CaCO 3 dissolution into the deep box (below 100 m depth). There are several other useful observations in the article by Milliman and Droxler (1996). They included an estimate of CaCO 3 burial, and they also estimated the input of alkalinity (expressed by them in terms of excess Ca input) from fluvial, groundwater, and hydrothermal sources. A major conclusion of their analysis was that CaCO 3 production, dissolution in the water column, and burial could be reconciled only if a majority of CaCO 3 dissolution occurs below surface waters but above the calcium carbonate saturation depth. These observations and their budget lead us to construct Fig. 4, as follows, for the contribution of CaCO 3 to the oceanic carbon pump. Net CaCO 3 production in surface waters (including both planktonic and benthic production) is estimated by Milliman and Droxler (1996) to total about mol year -1 (very close to the value estimated by Lee 2001). According to Milliman and Droxler (1996), about 2/3 of that surface production dissolves in the water column (deep water box), while the remainder is buried. A long-term geological return feed provides new DIC from fluvial, groundwater, and hydrothermal sources. CO 2 evasion from the surface water in response to the CaCO 3 production is readily calculated from the 0.6 rule (W = 0.6), to total about mol year -1. If we were to apply a constant value of 0.6 for W throughout the water column, then dissolving mol CaCO 3 year -1 would take up only mol CO 2 year -1 available from the oxidized organic matter. There would be net CO 2 release of mol year -1 (that is, mol year -1 CO 2 release in response to precipitation minus mol year -1 CO 2 uptake in response to dissolution). Fig. 4 Simple steady-state, 1-dimensional box model for the CaCO 3 portion of the oceanic inorganic carbon budget. The arrow for CO 2 evasion associated with precipitation is, of course, far smaller than the CO 2 invasion mediated by the production and downward particle flux associated with net organic production. The downward organic carbon fallout and oxidation in the water column account for the biologically driven CO 2 gas return feed. The geologically mediated alkalinity return feed is a slow cycle driven by chemical weathering and other processes of rock decomposition and returned to the ocean via surface run-off, groundwater, and hydrothermal fluids

9 Aquat Geochem (2011) 17: However, we can see from Figs. 2 and 3 that deep water W is substantially larger than that of the surface water throughout most of the Central Pacific Ocean. A value of 0.9 would balance the CO 2 evasion and invasion in the CaCO 3 cycle, while allowing for the external (geological timescale) cycling of DIC between PIC burial and return. OC cycling is also an important consideration. New production of OC in the photic zone totals * mol year -1 (Laws et al. 2000; Lee 2001). OC burial totals only about mol year -1, apparently largely of terrigenous origin (Hedges and Keil 1995); so we can consider OC surface-water production and deep water oxidation to be effectively balanced (the biological pump). These reactions take up atmospheric CO 2 into the surface ocean and vent it from the deep ocean with little OC accumulation in sediments. Because these reactions far exceed CO 2 fluxes associated with CaCO 3 reactions, the IC-generated CO 2 fluxes simply result in a slight diminution of the OC-generated CO 2 fluxes in the biological return feed shown on Fig. 4. Some value of w should be associated with CaCO 3 dissolution providing the geologically driven alkalinity return feed. Without some definition of the composition of the dissolving fluids, it is difficult to predict what the numerical value of this w term might be; it is likely to be rather variable. Whatever the average value might be, we believe that some equation like Eq. 3 dictates that there be (1-w) amount of rock-associated organic oxidation accompanying the CaCO 3 dissolution and compensating for the gas flux. 5 Conclusions The vertical distribution of w in the ocean provides insight into details of the functioning of the oceanic carbon cycle. Although W is ordinarily considered with respect to surface seawater equilibration with the overlying atmosphere, this coefficient is an intensive property, which also applies below the mixed layer. We have used the distribution of w at a single station in the North Pacific Ocean and along a transect through the Central Pacific (Figs. 2 and 3) to suggest a simple, but significant, modification of air-sea transfer for CO 2 gas in the oceanic IC balance (Fig. 4). The increase in w with depth and CaCO 3 dissolution there play the role, in the combined biological and carbonate pumps, of converting a larger fraction of OC-derived CO 2 back to HCO 3 than would occur from OC oxidation alone. It follows from the simple budgetary analysis performed with the assumption of steady state that the role of CaCO 3 reactions as a source or sink for atmospheric CO 2 depends on the total amount of CaCO 3 precipitated, the amount buried, and the difference in W between the sites of precipitation and dissolution. The assertion by Berner et al. (1983), that gas flux associated with CaCO 3 precipitation needs to be in balance, can be readily accepted in a stratified ocean. Of course the global ocean is neither 1-dimensional nor at steady state with respect to its carbon balance. The point of this calculation of w-induced deep ocean CO 2 flux from w calculated from a limited data set is not to provide a definitive budget of balancing CO 2 from CaCO 3 precipitation and dissolution. Rather, it is to demonstrate that the combination of CaCO 3 precipitation, partial CaCO 3 burial, a link with the OC cycle, and variable w through the water column can result in a balance (or a positive or negative imbalance) in air-sea gas transfer. The distributions of w in both surface waters and waters of maximum net OC oxidation between about 100 and 1,000 m are of great importance in evaluating oceanic CaCO 3 reactions as a net source or sink of atmospheric CO 2.

10 336 Aquat Geochem (2011) 17: OC reactions dominate the transfer of dissolved C from the mixed layer as particulate fallout downward in the water column. However, almost all of the OC decomposes in the water column rather than accumulating in the sediments. CaCO 3 burial far exceeds OC burial in transferring C from the water column to the sediments. It is therefore important to understand the net CO 2 transfer between the atmosphere and sediments associated with the CaCO 3 burial. We suggest that a more detailed evaluation of w throughout the global ocean would improve our understanding of the role of CaCO 3 reactions in the oceanic carbon balance. Acknowledgments We thank several people for their discussions of this problem over many years. We particularly identify the following supportive skeptics who challenged us to find a reconciliation between the 0.6 rule and the standard equation for CaCO 3 reactions: Bob Berner, Bob Garrels, Fred Mackenzie, and Roland Wollast. Fred Mackenzie, Bob Berner, Louis Legendre, and Dennis Swaney have all provided helpful comments on earlier drafts of this manuscript. The contribution of Michel Frankignoulle is gratefully acknowledged, particularly for his recognition of variation in the coefficient w. The manuscript has been greatly improved by the comments of two anonymous reviewers. We dedicate the paper to the memory of John Morse, a friend and respected colleague who was well known (among other things) for his interest in marine carbonate geochemistry. This work is a contribution to the European Project on Ocean Acidification (EPOCA), which receives funding from the European Community s Seventh Framework Programme under grant agreement References Berner RA, Maasch KA (1996) Chemical weathering and controls on atmospheric O 2 and CO 2 : Fundamental principles were enunciated by J. J. Ebelmen in Geochim Cosmochim Acta 60: Berner RA, Lasaga AC, Garrels RM (1983) The carbonate-silicate geochemical cycle and its effect on atmospheric carbon dioxide over the past 100 million years. Am J Sci 283: Boudreau BP, Middleburg JJ, Meysman FJR (2010) Carbonate compensation dynamics. Geophys Res Letters 37:L doi: /2009gl Broecker WS, Takahashi T (1966) Calcium carbonate precipitation on the Bahama Banks. J Geophys Res 71: Ebelmen JJ (1845) Sur les produits de la décomposition des espèces minérales de la famille des silicates. Ann Mines 7:3 66 Egleston, ES, Sabine, CL, Morel, FMM (2010) Revelle revisited: buffer factors that quantify the response of ocean chemistry to changes in DIC and alkalinity. Glob Biogeochem Cycles 24. doi: / 2008GB Feely RA, Sabine CL, Lee K, Millero FJ, Lamb MF, Greeley D, Bullister JL, Key RM, Peng T-H, Kozyr A, Ono T, Wong CS (2002) In situ calcium carbonate dissolution in the Pacific Ocean. Glob Biogeochem Cycles 16. doi: /2002GB Feely RA, Sabine CL, Lee K, Berelson W, Kleypas J, Fabry VJ, Millero FJ (2004) Impact of anthropogenic CO 2 on the CaCO 3 system in the oceans. Science 305: Frankignoulle M (1994) A complete set of buffer factors for acid/base CO 2 system in seawater. J Mar Syst 5: Frankignoulle M, Canon C, Gattuso J-P (1994) Marine calcification as a source of carbon dioxide: positive feedback of increasing atmospheric CO 2. Limnol Oceanogr 39: Hedges JI, Keil RG (1995) Sedimentary organic matter preservation: an assessment and speculative synthesis. Mar Chem 49: Lavigne H, Gattuso J-P (2010) Seacarb: seawater carbonate chemistry with R. R package version Laws EA, Falkowski PG, Smith WO Jr, Ducklow H, McCarthy JJ (2000) Temperature effects on export production in the open ocean. Glob Biogeochem Cycles 14: Lee K (2001) Global net community production estimated from the annual cycle of surface water total dissolved inorganic carbon. Limnol Oceanogr 46: Lerman A, Mackenzie FT (2005) CO 2 air-sea exchange due to calcium carbonate and organic matter storage, and its implications for the global carbon cycle. Aquat Geochem 11:

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