In situ calcium carbonate dissolution in the Pacific Ocean

3.1. Anthropogenic CO₂

[6] Our approach for estimating anthropogenic CO₂ in the Pacific Ocean is modified from the techniques described by Gruber et al. [1996]. Gruber et al. improved the earlier approaches of Brewer [1978] and Chen and Millero [1979] by developing the ΔC* method. This method is based on the premise that the anthropogenic CO₂ concentration (C_anth) can be isolated from measured DIC values (C_m) by subtracting the contribution of the biological pumps (ΔC_bio), the DIC waters that would have been in equilibrium with a preindustrial atmospheric CO₂ concentration of 280 ppm (C_eq280), and a term that corrects for the fact that surface waters are not always in equilibrium with the atmosphere (ΔC_diseq):

where

C_anth =

anthropogenic carbon concentration;

C_m

measured total carbon concentration;

ΔC_bio

change in DIC as a result of biological activity (both organic and inorganic);

C_eq280

DIC of waters in equilibrium with an atmospheric CO₂ concentration of 280 μatm;

ΔC_diseq

air-sea difference in CO₂ concentration expressed in μmol kg⁻¹ of DIC.

[7] The three terms to the right of the first equal sign make up ΔC*, which can be explicitly calculated for each sample. The fact that ΔC* is a quasiconservative tracer helps remedy some of the mixing concerns arising from the earlier techniques [Sabine et al., 1999; Sabine and Feely, 2001]. The ΔC_diseq term is evaluated over small isopycnal intervals using a water-mass age tracer such as CFCs [Sabine et al., 2002]. The quasiconservative tracer, ΔC*, is defined as the difference between the measured DIC concentration, corrected for biology and the concentration these waters would have at the surface in equilibrium with a preindustrial atmosphere (i.e., ΔC* = C_m − ΔC_bio − C_eq280). The ΔC* calculation used here is essentially the same as that originally defined by Gruber et al. [1996] with two small differences: a modification of the preformed alkalinity term based on the new global survey data and the addition of a denitrification term in the biological correction [see Sabine et al., 2002, for details].

3.2. Calcite and Aragonite Saturation

[8] The calcite and aragonite saturation levels were calculated using the program developed by Lewis and Wallace [1998]. The in situ degree of saturation of seawater with respect to calcite and aragonite is the ion product of the concentrations of calcium and carbonate ions, at the in situ temperature, salinity and pressure, divided by the stoichiometric solubility product for those conditions:

where the calcium concentration is estimated from the salinity, and the carbonate ion concentration is calculated from the TA data. The pressure effect on the solubility is estimated from the equation of Mucci [1983] that includes the adjustments to the constants recommended by Millero [1995].

3.3. CaCO₃ Dissolution

[9] The TA has three primary components in the water column: (1) the preformed TA (TA°), fixed when the parcel of water was last in contact with the atmosphere; (2) the changes in TA resulting from the release of protons during the oxidation of organic matter; and (3) the TA added to the water column via CaCO₃ dissolution. In this study, we employed the TA° equation of Sabine et al. [2002] based on near surface (0–60 m) data from the same Pacific data set used here:

where S is the salinity, PO = O₂ + R_O2/PHPO₄²⁻ [after Broecker, 1974], and θ is the potential temperature. The value of R_O2/P of 170 is used in this work [Anderson and Sarmiento, 1994].

[10] The organic matter correction of Chen [1978], based on changes in nitrate, was also adapted for use in this study. The largest component of this correction is from the oxidation of organic nitrogen. However, a coefficient of 0.63 was proposed by Kanamori and Ikegami [1982] to include contributions from organic phosphorus and sulfur as well. Rather than attempting to determine preformed nitrate values, Chen's equation was converted to apparent oxygen utilization (AOU) using the Anderson and Sarmiento [1994] N/O₂ ratio of 16/170. For a given isopycnal surface, the changes due to CaCO₃ dissolution can be evaluated using what we term TA*:

where NTA = (TA · 35)/S, NTA° = (TA° · 35)/S and AOU is the apparent oxygen utilization. Although this general approach has been used successfully for a number of years [e.g., Brewer et al., 1975; Chen, 1978; Chen et al., 1982; Chen, 1990; Sabine et al., 1995], the most common historical notation of ΔCa is somewhat misleading since this calculation does not involve calcium measurements and there are potential reactions, albeit minor, that could change this value without changing the dissolved calcium. We feel that the TA* term, in the spirit of the other star terms found in the recent literature, is more appropriate [Gruber et al., 1996].

4.1. Distributions of DIC and TA in the Pacific Ocean

[11] In the Pacific Ocean the lowest concentrations of DIC and TA are observed in surface waters. The surface concentrations (DIC range: 1975–2200 μmol kg⁻¹; TA range: 2200–2400 μmol kg⁻¹) are roughly correlated with salinity (Figures 2 and 3). DIC concentrations increase in the intermediate waters to form a large maximum (2325–2375 μmol kg⁻¹) at approximately 1800–2200 m in the North Pacific. In contrast, TA concentrations show local minima in the Antarctic Intermediate Water (AAIW) to the south and in the North Pacific Intermediate Water (NPIW) to the north. Below the North Pacific Intermediate Water, TA concentrations increase to a broad maximum at approximately 2200–4000 m (TA concentrations range from 2400–2460 μmol kg⁻¹). The general structure of the DIC and TA fields are similar to the density structure in the upper 1000 m. This similarity is an indication of the strong control that circulation plays in their distributions. The differences between the DIC and TA, particularly in shallow and intermediate waters, are caused by in situ remineralization processes. The DIC maximum is shallower than the TA maximum because DIC is more strongly influenced by the shallow remineralization of soft tissue organic carbon, whereas the TA is more strongly influenced by the dissolution of calcium carbonate particles deeper in the water column [Chen, 1990]. Except for the region north of 30°N, bottom waters have lower DIC and TA concentrations than the waters at midwater depths because the deep and bottom water circulation goes from south to north with upward mixing such that the oldest waters are centered at midwater depths (e.g., 2000–3000 m) of the North Pacific [Stuiver et al., 1983].

[12] The zonal DIC and TA isolines shoal from west to east along 30°N in the P2 section in the North Pacific (Figure 3). Deep ventilation near the Kuroshio Extension and the subsequent circulation in the subtropical gyre generates the zonal gradient of DIC and TA in the upper 1500 m of the water column. Both DIC and TA concentrations show the deepest ventilation near the coast and under the Kuroshio Extension west of 160°E. These results are consistent with the CFC and anthropogenic CO₂ distributions in the North Pacific that also indicate stronger ventilation processes in the western Pacific [Warner et al., 1996; Sabine et al., 2002].

4.2. Aragonite and Calcite Saturation Horizon Migrations

[13] The level at which aragonite and calcite are in thermodynamic equilibrium is the saturation depth. This depth is significantly shallower for aragonite than calcite because aragonite is more soluble in seawater than calcite (Figure 4). There is a pronounced shoaling of the aragonite and calcite from south to north and from west to east because of the higher DIC concentrations in northern and eastern regions relative to the TA concentrations (Figure 5; Feely et al., 1984). The general pattern of aragonite and calcite saturation is consistent with previous results [Takahashi, 1975; Chen et al., 1988; Feely and Chen, 1982; Feely et al., 1984, 1988; Kleypas et al., 1999].

[14] The anthropogenic CO₂ concentrations and the calculated aragonite and calcite saturation horizons are plotted for both present day (dashed line) and preindustrial levels (solid line) for the P15 and P2 sections (Figures 6a and 6b, respectively). The preindustrial levels are calculated by subtracting the anthropogenic CO₂ values [Sabine et al., 2002] from the DIC values presented here. When the preindustrial saturation horizons are compared to the present-day values, several distinct regions of upward migrations can be observed. The present-day aragonite saturation horizon ranges from approximately 200 m to 1320 m in the South Pacific with the greatest shoaling in the region from approximately 30°S to 5°S in the eastern South Pacific (Figure 4). In the North Pacific, the aragonite saturation horizon shoals to a minimum of 220 m at about 8°N, deepens to a maximum of about 580 m at 28°N, and shoals to its shallowest depth of approximately 120 m north of 50°N. An upward migration of the present-day saturation horizon relative to the preindustrial saturation horizon south of 38°S is observed to be between 30–80 m in the South Pacific and between 30–100 m in the North Pacific north of 3°N. A similar shoaling of the aragonite and calcite saturation horizon is observed in the east-west P2 section (Figure 6b). The west to east shoaling of the aragonite and calcite saturation horizons are consistent with the shoaling of the TA concentrations shown in Figure 3. This shoaling is the result of the deep ventilation and anticyclonic circulation in the North Pacific.

[15] The calcite saturation sections in Figures 6a and 6b indicate that the saturation horizon is close to 3000 m in the South Pacific and dramatically shoals to about 700 m just north of 20°N in the North Pacific. From there, it shoals to a minimum at approximately 250 m north of 50°N. The data suggest a distinct upward migration of the saturation horizon north of 20°N from the preindustrial period to the present ranging between 40 and 100 m. These results, indicating the shoaling of aragonite and calcite saturation horizons due to the effects of anthropogenic CO₂ ventilation in the surface and intermediate waters, imply that there may be a potential for enhanced dissolution of CaCO₃ particles in the undersaturated waters [Sarma et al., 2002].

4.3. CaCO₃ Dissolution Rates

[16] The WOCE/JGOFS global CO₂ survey in the Pacific Ocean was used for estimating CaCO₃ dissolution in the water column. Figure 7 shows a wedge-like cutout section of TA* along 170°W, 30°N, and 150°W. Positive concentrations of TA* are observed in the shallow waters near or slightly above the aragonite saturation horizon. Below this horizon, TA* concentrations increase rapidly from about 10–40 μmol kg⁻¹. This gradient is primarily located in the Intermediate Waters of the North and South Pacific. For example, north of 20°N, the largest increase occurs between about 400–1100 m, where the TA* increase is from <10 μmol kg⁻¹ in the North Pacific to values >40 μmol kg⁻¹ north of 40°N. In the South Pacific between 10°S and 45°S, where the aragonite and calcite saturation horizons are much deeper, the largest gradients are between 800 m to 1800 m. Farther south in the South Pacific, the TA* increase begins at much shallower depths, consistent with the shoaling of the aragonite saturation horizon in these waters. These results suggest that the extent of CaCO₃ dissolution is somehow related to the degree of aragonite saturation.

[17] Combining the above results with the apparent CFC-11 age data from the WOCE global survey data on isopycnal surface allows us to calculate the CaCO₃ dissolution rates by plotting TA* versus apparent CFC-11 ages (Figure 8). The slope of the line gives a value for the average CaCO₃ dissolution rate along the isopycnal surface. The method is limited to surface and intermediate waters where the apparent CFC ages are less than about 30–35 years due to mixing and dilution problems in the deeper waters. CaCO₃ dissolution rates calculated in this manner range from 0 in near-surface waters to a maximum exceeding 0.5 μmol kg⁻¹ yr⁻¹ in the intermediate waters (Figure 9). Lower rates are observed below the intermediate waters. The highest dissolution rates in the South Pacific were at depths between 1000 and 1800 m (σ_θ range: 27.3–27.6). In the North Pacific, the maximum dissolution rates were between 600 and 900 m (σ_θ range: 26.4–26.8).

[18] Since the CFC-age method is limited to water mass ages less than about 30–35 years, we used the natural ¹⁴C ages from samples collected on the same cruises to estimate CaCO₃ dissolution rates in deeper waters. The ¹⁴C decay rate of ∼1% every 83 years makes this isotope a good age tracer for dissolution processes in the deep sea. For waters >1500 m the mean CaCO₃ dissolution rate in the Pacific Deep Water was determined to be 0.051 μmol kg⁻¹ yr⁻¹, based on 853 data points. To define the spatial variability of the deep-water dissolution rates these data were subsetted by basin and examined on isopycnal surfaces in a manner similar to the shallow dissolution study. The dissolution rates were significantly lower than the shallow rates, ranging from 0.01 to 0.06 μmol kg⁻¹ yr⁻¹. The highest rates were observed at a σ_θ of 40.4 (Figure 10). In contrast to the shallow analysis, the South Pacific showed higher dissolution rates than the North Pacific. It is possible that the more corrosive waters of the North Pacific result in less carbonate particles actually reaching the deep waters. Contamination from bomb ¹⁴C prevents us from carrying the analysis to depths much shallower than ∼1000 m.

[19] One can get a sense of the dissolution throughout the water column from a plot of the average TA* as a function of depth (Figure 11). The maximum signal is near 1100 m, just below the average depth of the aragonite saturation horizon and well above the old southward return flow. A second local maximum is observed near 3800 m, just below the calcite saturation depth. A third maximum is observed in the bottom-waters. This could be a benthic dissolution signal. Although the waters at all levels intersect the bottom, by far the largest area of bottom is associated with waters deeper than 4500 m.