Effects of ocean acidification on the dissolution rates of reef-coral skeletons

Ocean acidification

As humans continue to burn fossil fuels at an unprecedented rate, the concentration of carbon dioxide (CO₂) in the atmosphere is presently higher than it has been for the last 420,000 years (Petit et al., 1999; Hansen et al., 2006; Hoegh-Guldberg et al., 2007). The oceans uptake a large proportion of the atmospheric CO₂, increasing the concentrations of both carbonic acid and bicarbonate ions, and reducing the concentration of carbonate ions, shifting the ocean’s acid–base balance toward a lower pH (Broecker, 1983; Caldeira & Wickett, 2003; Silverman et al., 2009). The increase in ocean acidification directly threatens calcifying marine organisms, such as reef-building corals and the myriad of species that rely on corals for protection and sustenance (Hoegh-Guldberg et al., 2007; Rodolfo-Metalpa et al., 2011).

Ocean pH has already decreased by 0.1 pH units since the 18th century, and is expected to drop by another 0.2–0.4 pH units by 2100. Yet the oceans are not homogenous in regard to rates of reductions in carbonate ions. Although warm waters increase reaction rates, thermodynamic principles and Henry’s Law tells us that cool temperate and polar waters absorb asymmetrically more CO₂ than tropical waters, and are therefore closer to unity than the more super-saturated tropical waters (Broecker, 1983). Yet the tropical oceans are changing at a more rapid rate and are acidifying more quickly than the cooler waters most likely because of the relationship of rapidly increasing ocean temperature and reaction rates (Zeebe et al., 2008). Moreover, the Pacific Ocean is more acidic than the Atlantic Ocean, and shoaling saturation depth is around 500 m in the Pacific and 4500 m in the Atlantic (Feely et al., 2004; Millero, 2007).

There is increasing evidence that ocean acidification, through the increase in the partial pressure of carbon dioxide (pCO₂) and the subsequent changes in the concentration of carbonate and bicarbonate ions, reduces rates of coral calcification, which are directly proportional to the saturation state of aragonite in the shallow oceans (Langdon & Atkinson, 2005). Other studies have shown that calcification rates are proportional to the concentration of carbonate ions in the water column (Anthony et al., 2008; Marubini et al., 2008). These studies are essentially synonymous, however, because the aragonite and calcite saturation state (Ω) is the product of the concentrations of calcium and carbonate ions divided by an equilibrium constant. Since the salt concentration, including calcium ions, stemming from terrestrial weathering hasn’t changed in the oceans for over 1.5 billion years, the aragonite saturation state is essentially a measure of carbonate ions in the oceans.

Perhaps more importantly is the strong interaction effects between temperature and ocean acidification on coral calcification rates (Reynaud et al., 2003; Erez et al., 2010). Indeed, the optimal window of physiological performance of a given marine species at a given temperature will be narrowed under acidification (Portner, 2010). Calcification of corals under ambient temperature do not necessarily change with increased pCO₂, but calcification decreases when both temperature and pCO₂ are elevated (Reynaud et al., 2003). Yet several studies have shown that many corals are unaffected by external carbonate ion concentrations because they have the capacity to up-regulate internal pH by actively exchanging internal hydrogen ions for calcium ions through Ca-ATPase transportation (Al-Horani, Al-Moghrabi & de Beer, 2003; Allemand et al., 2004; McCulloch et al., 2012). By modifying their internal chemistry, live corals may buffer themselves from ocean acidification. Coral skeletons, however, have no internal-buffering capacity because they are not protected by coral membranes (Rodolfo-Metalpa et al., 2011; Ries, 2011). Coral skeletons are instead subjected to the raw and immediate threats of ocean acidification and will be subjected to dissolution when the ocean’s pH declines.

Accretion of coral reefs

The accretion of coral reefs occurs over geological time periods when rates of calcium carbonate (CaCO₃) production exceed rates of destruction and dissolution (Neumann & MacIntyre, 1985; Buddemeier & Hopley, 1988; Glynn, 1997; Perry et al., 2013). The interaction between production and destruction depends on the consistency of coral cover through time. For example, where coral cover is consistently low, reef accretion is minimal (Neumann & MacIntyre, 1985). Most modern reefs, however, support little more than 28% live coral cover (Bruno & Selig, 2007), and are essentially veneers over pre-existing, antecedent foundations of CaCO₃ (Adey, 1978; Hopley, 1982). For example, the Florida Keys only supported, on average, 2–3% of live coral cover in 2011 (Office of National Marine Sanctuaries, 2011). Therefore, reefs with high carbonate cover and few live corals are particularly vulnerable to ocean acidification.

The average modern, shallow seaward coral reef in the Indo-Pacific, with high coral cover, has been estimated to produce about 4 kilograms of calcium carbonate per square meter of reef per year, which equates with an upward reef-growth rate of approximately 3 mm y⁻¹ (Smith & Kinsey, 1976). These estimates were based on alkalinity reduction techniques subjected to a pH of 8.2, equivalent to the pH of today’s oceans. By 2100 the ocean’s pH is expected to be 7.8, and we hypothesize that the destructive processes associated with ocean acidification might outweigh the constructive processes. The rates of dissolution of reef framework may, however, also depend on flow rates, the extent of cementation of reef framework, and on the porosity of corals and their surface area.

Reef cementation and coral porosity

Reefs vary in porosity depending on both: (i) the local rates of sedimentation and the extent to which those sediments become consolidated, or lithified, within the reef framework, and (ii) the extent of cementation of the reef framework. Both processes depend in part on exposure to water-flow rates (MacIntyre & Marshall, 1988). High-energy, windward reefs consistently exposed to large waves are generally more highly cemented than low-wave energy, leeward reefs because mass-transfer rates influence rates of cementation. Cementation involves the infilling of intra-skeletal pores with either Mg calcite or aragonite (MacIntyre & Marshall, 1988). While the extent of cementation affects the dislodgment of reef substrate and the tenacity of corals to remain attached to reefs during storms (Madin, Hughes & Connolly, 2012), the extent of reef cementation may also affect dissolution rates during ocean acidification because the infilling of pores by cements decreases the surface area of exposure (Cubillas et al., 2005).

Reef corals also vary in porosity (Gladfelter, 1982; Hughes, 1987). Although all modern corals secrete orthorhombic aragonite fusiform crystals, as small as 1–3 µm (Gladfelter, 1982), corals vary considerably in the arrangement of the crystals, which influences the internal surface area that is exposed (Fig. 1). Fast-growing corals, such as Montipora and Acropora, are mostly perforate corals (Gladfelter, 1982), whereas slow-growing corals, such as Pectinia and Symphyllia, are imperforate (Table 1). An extreme example of imperforate skeletons is evident by the observation of occasional floating, massive Symphyllia colonies (DeVantier, 1992). Because of the fused nature of the dissepiments and their imperforate skeletons, gases are trapped in the septal chambers and upon dislodgement from reefs, for example during a storm, the colonies will float. Perforate corals, however, do not have the capacity to isolate septal chambers.

The internal porosity of coral skeletons, at the scale of 0.5–1 mm (Fig. 1), increases the available surface area of chemical exchange and therefore increases the potential rates of dissolution. Walter & Morse (1984) showed that rates of dissolution of skeletal carbonates were inversely related to grain-size diameter and surface roughness, with fine grained carbonates dissolving faster than large, rough surfaces. However, we should not discount the possibility that perforate and imperforate corals also differ in other aspects, beyond the obvious differences in porosity, and therefore we question whether surface area is a useful predictor of rates of passive dissolution of both perforate and imperforate corals.

This study will examine whether the porosity and the surface area of coral skeletons will influence their rate of dissolution when the ocean pH is 7.8, which is predicted to occur by 2100. More specifically, we tested three hypotheses: (1) that perforate Montipora coral skeletons are more likely to passively dissolve than imperforate Pectinia coral skeletons at a pH of 7.8, (2) that the rates of passive dissolution of coral-colony skeletons are proportional to their surface areas, and (3) future reef accretion rates under ocean acidification will differ depending on the nature of the coral assemblages, with perforate coral assemblages unable to keep up with predicted sea-level rise and imperforate coral assemblages faring a better chance at keeping up with sea-level rise and ocean acidification.

Acidification experiments

In order to test the first hypothesis, perforate Montipora colonies (Fig. 1) and imperforate Pectinia colonies without tissue (Fig. 2) were used to make comparisons of weight loss when immersed in seawater and held in zero-flow conditions (i.e., to test passive dissolution) at a pH of 8.2, equivalent to the pH of today’s oceans, and compared with colonies held at a pH of 7.8, which is predicted to occur globally by 2100. Fifteen skeletal samples (≤5 cm) of Montipora spp. colonies and fifteen skeletal samples of Pectinia spp. were collected from the fringing reefs of Okinawa, Japan in 2001. In order to test the second hypothesis, we used a variety of growth forms of Montipora, including submassive, branching, encrusting, and foliose. Colonies of Pectinia with different surface areas were used for experimental treatments, but all samples were foliose because Pectinia is only found as foliose colonies on modern coral reefs.

Before pH treatments, the samples were placed in a drying oven at 40°C for 48 h and weighed (g) using a Sartorius Research Balance. Each treatment sample was then placed in a separate container of seawater that was maintained at a pH of 7.8 by adding diluted acetic acid to match the predicted pH of the seawater in the year 2100 (Intergovernmental Panel on Climate Change (IPCC, 2007)). The control samples were placed in seawater that was maintained at a pH of 8.2, to match modern ocean conditions, and maintained at 24°C and a salinity of 35. Total alkalinity was not measured in this study. Seawater was changed every 2 days. After 7 days the samples were rinsed and dried in a drying oven at 40°C for 48 h, and re-weighed. The volume of each coral sample (mL) was calculated using a displacement method and the surface area of each coral sample (cm²) was calculated using a single wax-dipping method (Veal, Carmi & Fine, 2010).

Data analyses

The difference in dry weight (g) before and after the acid treatment was calculated for each coral sample. To correct for differences in initial weight, the loss of calcium carbonate was divided by each coral’s initial weight. To compare differences in dissolution rates that may have varied in accordance with growth form, we undertook an analysis of variance (ANOVA) and a Tukey’s post-hoc test using R (R Development Core Team, 2012). The relationship between the surface area, volume, and the loss of calcium carbonate was examined using curve fitting with Matlab^®.

Accretion-dissolution model

The loss of calcium carbonate was extrapolated from the change in calcium carbonate per gram cm⁻² d⁻¹, to the equivalent loss of calcium carbonate per kg m⁻² y⁻¹. This loss was compared with the geological literature and converted to the approximate equivalent of vertical reduction of reef framework in mm per year (Smith & Kinsey, 1976). The loss was compared with predicted sea-level rise (Vermeer & Rahmstorf, 2009). In order to achieve this goal, the reef accretion rates were modeled as an ordinary differential equation: (1)${\displaystyle \mathrm{d}A/\mathrm{d}t=(a.\mathrm{A})/\mathrm{A}+b.\mathrm{S}-(c.\mathrm{D})/\mathrm{A},}$ where A is the accretion of a reef relative to time (t); a is the accretion coefficient determined by coral and coralline algal growth minus the bioerosion rates (input as 7 mm y⁻¹ for reefs that accrete the maximum of 10 kg CaCO₃ m⁻² y⁻¹; 3 mm y⁻¹ for reefs that accrete 4 kg CaCO₃ m⁻² y⁻¹; and 0.75 mm y⁻¹ for reefs that accrete 1 kg CaCO₃ m⁻² y⁻¹, with a 50% average reef porosity, after Kinsey, 1979; Smith, 1983); b is a coefficient for sedimentation (S), input as 1 mm per year for consistency; and c is a coefficient for the dissolution (D) rates. The equations were solved using Runge–Kutta methods using the ode45 solver in Matlab^® (code is available in the Appendix S1).

The results of passive dissolution were input into our reef-growth model and compared with projections of global sea-level rise, from 1990 to 2100 following Vermeer & Rahmstorf (2009), which did not consider regional isostatic rebound effects, regional tectonics, and local land-use effects. The sea-level rise projections used different IPCC (2007) emission scenarios, including the B1 scenario representing a +1.8°C global increase in temperature, the A2 scenario representing a +3.4°C global increase in temperature, and the A1F1 scenario representing a 4°C global increase in temperature.

Accretion-dissolution model

The sea-level rise projections from 1990 to 2100 were constructed using different IPCC (2007) emission scenarios, including the B1 scenario, representing a +1.8°C global increase in temperature, the A2 scenario representing a +3.4°C global increase in temperature, and the A1F1 scenario representing a 4°C global increase in temperature (Fig. 6; Vermeer & Rahmstorf, 2009). These sea-level projections were compared with three different reef-building capacities in conjunction with rates of perforate and imperforate coral skeletons (Eq. (1)) under ocean acidification (Figs. 6 and 7). The modeled reef with high dissolution rates, which included perforate skeletons, and consistently high coral cover (10 kg CaCO₃ m⁻² y⁻¹) is not expected to keep up with sea level rise under ocean acidification (Fig. 6). By contrast, the modeled reef with low dissolution rates, which included imperforate corals, is expected to continue to grow reefs and keep up with sea level rise, but only reefs consistently supporting high coral cover (10 kg CaCO₃ m⁻² y⁻¹) and only to 2050. Around 2050, the model shows that the rates of sea level rise are expected to increase faster than the rates at which corals can grow reefs (Fig. 7).