*Article* **Impacts of Warming and Acidification on Coral Calcification Linked to Photosymbiont Loss and Deregulation of Calcifying Fluid pH**

**Louise P. Cameron 1,2, Claire E. Reymond 2,3, Jelle Bijma 4, Janina V. Büscher 5, Dirk De Beer 6, Maxence Guillermic 7, Robert A. Eagle 7, John Gunnell 1, Fiona Müller-Lundin 1, Gertraud M. Schmidt-Grieb 5, Isaac Westfield 1, Hildegard Westphal 2,8,9 and Justin B. Ries 1,2,\***


**Abstract:** Corals are globally important calcifiers that exhibit complex responses to anthropogenic warming and acidification. Although coral calcification is supported by high seawater pH, photosynthesis by the algal symbionts of zooxanthellate corals can be promoted by elevated pCO2. To investigate the mechanisms underlying corals' complex responses to global change, three species of tropical zooxanthellate corals (*Stylophora pistillata*, *Pocillopora damicornis*, and *Seriatopora hystrix*) and one species of asymbiotic cold-water coral (*Desmophyllum pertusum*, syn. *Lophelia pertusa*) were cultured under a range of ocean acidification and warming scenarios. Under control temperatures, all tropical species exhibited increased calcification rates in response to increasing pCO2. However, the tropical species' response to increasing pCO2 flattened when they lost symbionts (i.e., bleached) under the high-temperature treatments—suggesting that the loss of symbionts neutralized the benefit of increased pCO2 on calcification rate. Notably, the cold-water species that lacks symbionts exhibited a negative calcification response to increasing pCO2, although this negative response was partially ameliorated under elevated temperature. All four species elevated their calcifying fluid pH relative to seawater pH under all pCO2 treatments, and the magnitude of this offset (Δ[H+]) increased with increasing pCO2. Furthermore, calcifying fluid pH decreased along with symbiont abundance under thermal stress for the one species in which calcifying fluid pH was measured under both temperature treatments. This observation suggests a mechanistic link between photosymbiont loss ('bleaching') and impairment of zooxanthellate corals' ability to elevate calcifying fluid pH in support of calcification under heat stress. This study supports the assertion that thermally induced loss of photosymbionts impairs tropical zooxanthellate corals' ability to cope with CO2-induced ocean acidification.

**Keywords:** microelectrode; ocean acidification; global warming; calcifying fluid; scleractinian coral; zooxanthellate photosymbiont; photosynthesis; calcification; bleaching

**Citation:** Cameron, L.P.; Reymond, C.E.; Bijma, J.; Büscher, J.V.; De Beer, D.; Guillermic, M.; Eagle, R.A.; Gunnell, J.; Müller-Lundin, F.; Schmidt-Grieb, G.M.; et al. Impacts of Warming and Acidification on Coral Calcification Linked to Photosymbiont Loss and Deregulation of Calcifying Fluid pH. *J. Mar. Sci. Eng.* **2022**, *10*, 1106. https://doi.org/10.3390/ jmse10081106

Academic Editor: Weidong Zhai

Received: 20 July 2022 Accepted: 6 August 2022 Published: 12 August 2022

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**Copyright:** © 2022 by the authors. Licensee MDPI, Basel, Switzerland. This article is an open access article distributed under the terms and conditions of the Creative Commons Attribution (CC BY) license (https:// creativecommons.org/licenses/by/ 4.0/).

#### **1. Introduction**

Anthropogenic emissions are predicted to cause sea-surface warming [1] and ocean acidification (OA)—a process that lowers seawater pH and aragonite saturation state [2] (ΩA). OA increases both the dissolution rate of CaCO3 shell/skeleton [3] and the rate at which new shell/skeleton is formed [4]. Tropical scleractinian corals are carbonate producers [5] that acquire nourishment via symbiotic photosynthetic zooxanthellae and from heterotrophic feeding [6]. They are vulnerable to warming as many exist near the upper end of their thermal tolerance limits [7]. Warming beyond a coral's thermal tolerance may cause the loss of photosymbionts—a process known as bleaching.

Coral calcification occurs in the coral's calcifying fluid, which is influenced by both external seawater chemistry and the coral itself [8]. Corals elevate saturation state of this fluid through a combination of pH elevation via the active removal of protons using membrane-bound Ca2+-ATPase proton pumps [8–14] and DIC elevation [15,16]. Photosynthesis may aid pH elevation by supplying the necessary ATP required to drive Ca2+-ATPase proton pumps. Under ocean acidification, the formation of CaCO3 is more energetically costly [11,17,18], although this energetic cost is negligible compared to the total proportion of energy produced through photosynthesis [8]. Photosymbionts may therefore play a crucial role in mitigating the impacts of OA on corals by providing enough ATP to increase ion pumping rates in support of calcification. Photosynthesis may also be responsible for metabolic DIC elevation at the site of calcification, as implied by long-term seasonal variations in the concentration of DIC in the calcifying fluid [16]. Under warming, bleaching may impair coral calcification by reducing the amount of photosynthate that is translocated to the coral host, thereby increasing the coral's reliance on heterotrophic feeding [6], decreasing the metabolically derived DIC pool, and, potentially, reducing the amount of ATP available for proton-regulation in support of calcification.

Although tropical scleractinian corals generally exhibit parabolic calcification responses to ocean warming that are centered on their thermal optima [19], their calcification responses to OA are more nuanced. Many species exhibit reduced calcification rates [20,21], while others exhibit threshold responses [22], parabolic responses [19,21], or no calcification response to acidification [13,23,24].

This highlights the complexity of coral biomineralization, which can be biologically mediated by the secretion of skeletal organic molecules (SOM) such as adhesion, signaling, and structural proteins (e.g., calmodulin and sulphated acidic proteoglycans) [25–27]. The abundance of these SOMs in the calcifying fluid has been shown to alter the rate and morphology of aragonite precipitated [28–32].

Cold-water corals that inhabit deeper aphotic environments generally lack zooxanthellae and acquire nourishment exclusively through heterotrophic feeding [33]. The systems for regulating calcifying fluid pH (pHCF) within azooxanthellate corals may therefore differ from zooxanthellate corals. Azooxanthellate corals exhibit reduced skeletal density [34], reduced calcification rate [35], and altered rates of respiration and feeding [36] in response to OA, but increased calcification rates under elevated temperature [35,37].

Numerous studies have investigated the isolated effects of ocean acidification (e.g., [11,24,38,39]) and thermal stress (e.g., [40]) on coral pHCF. However, the present study—along with its companion paper [41]—are the first to investigate both the independent and combined effects of warming and acidification on coral pHCF. These empirical constraints, combined with photosymbiont and calcification data for the three tropical coral species, yield insight into the mechanism by which warming and acidification interact to so negatively impact coral growth. Examining the roles that pHCF regulation and symbiont abundance play in the coral calcification response to OA and warming should improve understanding and prediction of how different species of corals will respond to future global change.

We investigate these relationships by culturing three species of tropical zooxanthellate corals (*Stylophora pistillata*, *Pocillopora damicornis*, *Seriatopora hystrix*) and one azooxanthellate cold-water species (*Desmophyllum pertusum*, syn. *Lophelia pertusa*) under control

(tropical = 28 ◦C, cold-water = 9 ◦C) and elevated (tropical = 31 ◦C, cold-water = 12 ◦C) temperatures at near-present-day (451–499 ppm), year 2100 (885–1096 ppm), and year 2500 (2807–3194 ppm) pCO2 scenarios [1]. Coral calcifying fluid pH was measured with protonsensitive microelectrodes. Symbiont abundance of the zooxanthellate corals was estimated based on coral color to investigate the role of photosymbionts in the coral calcification response to warming and acidification.

This work is a companion paper to Guillermic et al. [41], which investigated the impacts of pCO2 and temperature on the calcifying fluid chemistry and calcification rate of two tropical species of scleractinian corals, *S. pistillata* and *P. damicornis.* The present study builds upon its companion paper by investigating these relationships for two additional species of scleractinian corals—the tropical species *S. hystrix* and the cold-water species *L. pertusa.* The present study also includes photosymbiont data for all three tropical coral species cultured under all experimental treatments.

#### **2. Materials and Methods**

#### *2.1. Overview of Experimental Design*

Three species of tropical scleractinian corals (*S. pistillata*, *P. damicornis*, and *S. hystrix*) were cultured under three ocean acidification (OA) scenarios established by modification of pCO2 at control and elevated temperatures (control acidification: 28.29 ◦C ± 0.01 s.e./466 ppm ± 8; 31.72 ◦C ± 0.06/499 ppm ± 9; moderate acidification: 27.88 ◦C ± 0.02/925 ppm ± 15; 30.83 ◦C ± 0.02/885 ppm ± 12; high acidification: 28.17 ◦C ± 0.02/2807 ppm ± 119; 30.93 ◦C ± 0.02/3194 ppm ± 135) in four replicate tanks at each treatment for 62 days in April–June 2016 (Table 1). The control temperature treatment was assigned to fall within the natural temperature range of coral reefs in Fiji (26–29 ◦C), where these corals were collected [42]. Likewise, the elevated temperature was assigned to be just above the bleaching threshold (30–30.5 ◦C) for corals in Fiji [42]. Simultaneously, the cold-water coral *L. pertusa* was cultured under three OA scenarios at control and elevated temperatures (control acidification: 8.86 ◦C ± 0.02/451 ppm ± 24; 12.17 ◦C ± 0.03/494 ppm ± 16; moderate acidification: 9.07 ◦C ± 0.03/1096 ppm ± 99; 12.52 ◦C ± 0.04/1079 ppm ± 60; high acidification: 8.83 ◦C ± 0.02/2864 ppm ± 222; 12.67 ◦C ± 0.03/3167 ppm ± 202) in four replicate tanks at each treatment for 33 days in April–May 2016 (Table 2). Corals were acclimated to laboratory conditions for one week and then to experimental conditions for an additional two weeks prior to the start of the experiment. Calcification rates, pHCF, and relative photosymbiont abundance were quantified during the experiment (details below).

#### *2.2. Coral Husbandry*

Experiments were carried out in the MAREE marine experimental facility at the Leibniz Centre for Tropical Marine Research (ZMT). Specimens of the tropical scleractinian coral species *S. pistillata* and *P. damicornis* were obtained from DeJong MarineLife (Netherlands). Colony-level information was not available for these specimens. Specimens of the tropical scleractinian coral species *S. hystrix* were obtained from an experimental stock colony provided by the ZMT. Fragments of the cold-water coral *L. pertusa* were obtained from four colonies of an experimental stock culture provided by the Marine Biogeochemistry Department of the Helmholtz Centre for Ocean Research in Kiel (GEOMAR), previously collected at a depth of ca. 200 m at Nord-Leksa Reef in Trondheimsfjord, Norway. Coral fragments (*S. pistillata* = 65, *P. damicornis* = 63, *S. hystrix* = 52, and *L. pertusa* = 43) were mounted onto 3 cm x 3 cm plastic egg-crate stands using cyanoacrylate epoxy and assigned a unique identifier. Each treatment was replicated in four tanks. Equivalent size ranges of specimens were maintained across treatments. Coral specimens obtained from larger colonies were randomly distributed amongst pCO2 and temperature treatments and replicate tanks. Coral specimens that died before the completion of the experiment were promptly removed from the tanks so as not to impact the remaining live corals in the experiment. Additional details about the number and weight of individuals in each treatment are provided in Section S1 of the Supplementary Online Material.

**Table 1.** Average calculated parameters for the tropical corals and all treatments: pCO2 of the mixed gases in equilibrium with seawaters (pCO2 (gas-e)), pH on seawater scale (pHSW), carbonate ion concentration ([CO3 <sup>2</sup>−]), bicarbonate ion concentration ([HCO3 −]), dissolved carbon dioxide ([CO2]SW), and aragonite saturation state (ΩA). Average measured parameters for all treatments: salinity (Sal), temperature (Temp), pH on NBS scale (pHNBS), total alkalinity (TA), and dissolved inorganic carbon (DIC). 'SE' represents standard error and 'n' is the sample size.



**Table 2.** Average calculated parameters for *Lophelia pertusa* and all treatments: pCO2 of the mixed gases in equilibrium with seawaters (pCO2 (gas-e)), pH on seawater scale (pHSW), carbonate ion concentration ([CO3 <sup>2</sup>−]), bicarbonate ion concentration ([HCO3 −]), dissolved carbon dioxide ([CO2]SW), and aragonite saturation state (ΩA). Average measured parameters for all treatments: salinity (Sal), temperature (Temp), pH on NBS scale (pHNBS), total alkalinity (TA), and dissolved inorganic carbon

After 1-week of acclimation at control conditions, temperature and pCO2 were then incrementally increased to treatment levels over an additional week, after which time coral specimens were acclimated to treatment conditions for an additional week prior to the experiment. All coral specimens were cultured in 10-L replicate tanks supplied with seawater from 244-L sumps, where water filtration, temperature control, and pCO2 control occurred. Seawater was filtered with protein skimmers, mechanical filters, and activated charcoal. All tropical coral aquaria were illuminated with 150 lux using actinic blue and white aquarium lights on a 12-h light/dark cycle. Aquaria holding the cold-water coral *L. pertusa* were not illuminated as this species lives below the photic zone. Each experimental treatment containing *L. pertusa* specimens (comprised of 4 replicate tanks) shared a water source with a separate reservoir containing five specimens of the king scallop *Pecten maximus* cultured as part of a separate experiment [43]. Seawater was filtered with protein skimmers, mechanical filters, and activated charcoal before returning to the *L. pertusa* tanks.

During acclimation and experimental periods, the tropical corals were fed 1-day old *Artemia salina* nauplii hatched from ca. 40 mg of eggs. Approximately 10 mL of concentrated live nauplii were introduced into each replicate tank every second day. The food mixture was pipetted directly adjacent to each coral specimen. Specimens of *L. pertusa* were fed the same diet supplemented with 20 mL of *Calanus finmarchicus* concentrate (Goldpods) suspended in the initial aliquot of *Artemia salina*. Corals were fed at the end of the day, and all filtration material was cleaned the following morning.

#### *2.3. Seawater Chemistry Manipulation and Measurement*

Measured and calculated carbonate system parameters are summarized in Tables 1 and 2. Experimental tank pCO2 was maintained by vigorously bubbling mixtures of CO2-free air and CO2 into the 244-L treatment sumps with microporous sparging tubes. The pCO2 of the bubbled gases was achieved by mixing compressed CO2-free air and compressed CO2 with solenoid-valve mass flow controllers at flow rates proportional to the target pCO2 conditions. Natural seawater, originally collected from Spitsbergen, Norway, was continuously added to each of the 244-L sumps at a rate of 0.6 L/hour. Temperature (s.e.) for the tropical corals was maintained at 28 (0.02) ◦C and 31 (0.06) ◦C using 125-watt aquarium heaters (EHEIM), controlled with a programmable thermostat. Temperature for *L. pertusa* was maintained at 9 ◦C (0.03) and 12 ◦C (0.04) using aquarium chillers (Aqua Medic).

Temperature, pH, and salinity of all replicate tanks were measured three times per week using a multi-electrode probe (WTW Multi 3430 Set K). Samples for the analysis of dissolved inorganic carbon (DIC) and total alkalinity (TA) were collected weekly from each of the replicate tanks at midday and used to calculate other carbonate system parameters using the program CO2SYS [44]. Nutrient concentrations ([NO3 −], [PO4 <sup>3</sup>−], and [NH4 +]) of all replicate tanks were measured weekly. Additional details about the methods used to measure the carbonate system and nutrient concentrations of replicate tanks are provided in Section S2 of the Supplementary Online Materials.

#### *2.4. Calcification Rates*

Calcification rates were calculated from the change in estimated dry weight of all coral fragments over the experimental period. Dry weights were estimated from buoyant weight measurements taken at the start and end of the experimental period according to the following empirically derived relationships:

*Stylophora pistillata*: Dry weight (g) = 1.919 × Buoyant Weight (g) + 7.677; *Pocillopora damicornis*: Dry weight (g) = 1.662 × Buoyant Weight (g) + 8.777; *Seriatopora hystrix*: Dry weight (g) = 1.668 × Buoyant Weight (g) + 8.493; *Lophelia pertusa*: Dry weight (g) = 1.594 × Buoyant Weight (g) − 0.206;

where the precision of this relationship is equivalent to the standard error of the regression (*S. pistillata*: 0.060 g; *P. damicornis*: 0.053 g; *S. hystrix*: 0.068 g; *L. pertusa*: 0.023 g). Additional details about the methods used to calculate calcification rate via buoyant weights are provided in Section S3 of the Supplementary Online Materials.

The number of days between the start and end buoyant weight measurements was then used to standardize %-calcification to a daily rate. Coral skeletons were also labeled with the fluorescent dye calcein (30 mg Se-Mark liquid calcein/kg-seawater) for 5 days prior to the initial buoyant weighing to identify skeletal material produced exclusively under the experimental conditions. Although all four species of corals recorded the calcein marker in their coral skeleton, rates of linear extension could not be reliably measured from the calcein marker because the dye was not incorporated into the skeletons in a consistent manner (see Section S4 of the Supplementary Online Materials for images of coral uptake of the calcein dye).

#### *2.5. Estimating Coral Photosymbiont Index*

The tropical coral specimens were photographed alongside the Coral Watch Coral Health Chart color scale [45–47] (Section S5 of the Supplementary Online Materials) under 150 lux (i.e., equivalent lighting to their experimental treatments) at the end of the experimental period to estimate relative photosymbiont abundance (a proxy for bleaching) of the coral specimens. This method involved extracting red-band color of the live coral tissue and the color scale and then assigning each pixel within the coral tissue image a discrete score (1–5) relative to the red-band values of the color scale (Figure 1). Additional details about the methods used to process photographs used for the estimation of photosymbiont index are provided in Section S5 of the Supplementary Online Materials.

**Figure 1.** Representative images used in the estimation of relative photosymbiont abundance. Panel (**A**) depicts a healthy, unbleached coral from the control temperature, high pCO2 treatment (color score = 5.69). Panel (**B**) depicts a partially bleached coral from the high temperature, high pCO2 treatment (color score = 4.56).

#### *2.6. Measurement of Calcifying Fluid pH*

Calcifying fluid pH was measured using proton-sensitive liquid ion-exchanger (LIX) microelectrodes produced at the Max Planck Institute for Marine Microbiology (MPIMM) using a modified version of the technique described in De Beer et al. [48]. In brief, green soda lime glass microcapillary tubes (Schott model 8516) were held in a heated coil and pulled to a target tip diameter of ca. 10 μm, yielding final diameters of 8–20 μm. These were then silanized to produce a hydrophobic surface that allowed the adhesion of the LIX membrane. The microcapillary tubes were filled with ca. 300 μm of degassed, filtered electrolyte (300 mM KCl, 50 mM sodium phosphate adjusted to pH 7.0) using a plastic syringe with a 0.1-mm tip. The microcapillary tubes were then backfilled with LIX containing a polyvinyl chloride (PVC) epoxy to prevent leakage of electrolyte by submerging the tips of the microcapillary tubes in LIX and apply suction to the other end of the tube until the PVC-containing LIX was drawn into the tip of the microcapillary by 100–200 μm. Microcapillary tubes were encased in a Pasteur pipette for shielding, with the pulled tip of the microcapillary tube protruding ca. 2 cm beyond this casing. This casing was filled with a 0.3 M KCl solution and connected to the reference electrode with an Ag/AgCl wire to minimize electrical noise. Microelectrodes were left for 24 h after construction to allow for stabilization of the LIX membranes.

All microelectrode equipment (millivolt meter, National Instruments DAQ Pad 6020E, laptop, cables, micromanipulator, VT80 Micos motor arm, lab stands, Zeiss Stemi SV6 binocular microscope) was set up adjacent to the experimental tanks to minimize transport stress for the corals. Two reservoirs of seawater, sourced from the corresponding experimental treatment tanks, were established next to the microelectrode system. These reservoirs were bubbled with the corresponding treatment gases and maintained at the corresponding treatment temperature using aquarium heaters or chillers. The seawater was circulated between the two reservoirs through two 5.4 L flow-through chambers (30 × 12 × 15 cm). All pH microelectrode measurements were performed within these smaller flow-through chambers.

Measurements of calcifying fluid pH were made in the flow-through chambers filled with treatment seawater. Light levels in these chambers were measured using a digital lux-meter positioned next to the target coral polyp and were maintained at 150 lux. All corals were acclimated to the microelectrode chamber until polyp extension was observed prior to measurements (minimum of 10 min). Measurements of calcifying fluid pH were performed on three replicate individuals per treatment, with one measurement obtained for each individual. Calcifying fluid pH measurements were obtained for all species in all pCO2 treatments under the control temperature treatment. Due to constraints on time and resources available for the experiment, calcifying fluid pH measurements under the high temperature treatment were only obtained for one species (*S. pistillata*).

The proton-sensitive LIX microelectrodes were used to measure both seawater and calcifying fluid pH. Before and after measurement of calcifying fluid pH, all microelectrodes were calibrated at the treatment temperature with pH 7 and 9 NBS buffers. The vertical position of the microelectrode was controlled with one-micron precision using a motorized micromanipulator. The microelectrodes were slowly inserted with a micromanipulator through the coral tissue into the upper portion of the coral calyx, between septal ridges and proximal to the thecal wall, until the skeleton was reached. This positioning of the electrode was verified by a shift in the pH profile [14] (Figure 2). A vertical pH profile (Figure 2) was then obtained by moving the microelectrode out of the calyx into the adjacent seawater. This profile was obtained in 1 μm steps for the first 20 μm, followed by 5 μm steps out into the surrounding seawater.

**Figure 2.** An example of a vertical pH profile generated during measurement of the coral calcifying fluid pH. This particular pH profile was generated for a specimen of *S. hystrix* cultured in the '1000 ppm pCO2, 28 ◦C' treatment.

The 1-μm spatial resolution of the micromanipulator allowed for the positioning of the electrode within the thin calcifying fluid immediately adjacent to the coral skeleton. If the skeleton was inadvertently contacted during this positioning, it is possible that the tip of the microelectrode would break and render it dysfunctional. It was visually evident if the microelectrode tip broke upon contact with the skeleton, and this would also result in an abrupt voltage anomaly, often followed by a drift in the voltage even while the electrode was in a fixed position. The pH profile was aborted if there was evidence of any of these issues and then reinitiated with a new microelectrode.

The calibration and microelectrode pH data were processed by parsing scatter-plots of the data into three zones, which were annotated at the time of data collection. The calibration data were parsed as pH 7 buffer and pH 9 buffer. The microelectrode pH data were parsed as calcifying fluid, tissue, and seawater. Notes recorded during the original measurements were used to assist in identifying boundaries of adjacent zones. Measured mV within each zone of the calcifying fluid measurements were converted to pH using the calibration regression produced for each microelectrode. The Δ[H+] was calculated for each measured coral as the difference between the proton concentration ([H+]) of the coral's surrounding seawater and the [H+] of the coral's calcifying fluid, both measured with the calibrated, proton-sensitive LIX microelectrodes.

#### *2.7. Statistical Analysis*

Statistical analyses were carried out in *R*. Corals that did not survive the experimental period (see Section S1 of the Supplementary Online Material) were excluded from analyses. A series of linear mixed effects models (lmers) were performed to investigate the influence of seawater pCO2 and temperature on coral physiology (calcification rate, calcifying fluid pH, Δ[H+], photosymbiont index), with treatment tank specified as a blocking factor [49]. Akaike information criterion (AIC) was used to estimate the relative amount of information lost by any given model [50]. The final model was chosen based on the lowest AIC score (whereby a lower score reflects a better fitting model) and highest R2, which reflects the goodness of fit (from 0 to 1, 1 being a perfect fit) (see Section S6 of the Supplementary Online Material for AIC model selection tables). The normality and homoscedasticity of the chosen tests were then analyzed using diagnostic plots (QQ-plot, residuals vs. fitted plot), and normality was tested using a Shapiro–Wilk test. Color scores were square-root transformed to meet the assumption of normality. If an interaction term was significant, the individual levels of that interaction were examined in their own linear mixed effects models to interpret main effects.

Analysis of co-variance (ANCOVA) was used to investigate the impacts of seawater pCO2, temperature, and species on calcification rate, thereby allowing interspecific comparisons of the impacts of OA and warming on coral calcification rate. An ANCOVA was also used to make interspecific comparisons of the impact of OA on calcifying fluid pH and Δ[H+] of different coral species. The latter analyses excluded individuals from the elevated temperature treatment, as calcifying fluid pH was only obtained for one of the three coral species in this treatment.

Linear mixed effects models were used to examine the impact of photosymbiont index, as a proxy for photosymbiont abundance, on calcification rate. Multiple linear models were generated, starting with the model that contained the most terms (i.e., modeling calcification rate as a function of seawater pCO2, temperature, and photosymbiont index). Final interpretations of the data were based upon the models that maximized R2 and minimized AIC (see Table S8). An alpha of 0.05 was used for all models, whereby any relationship with a *p*-value of <0.05 was deemed statistically significant.

#### **3. Results**

#### *3.1. Predictors of Calcification Rate*

The calcification rate of all coral species (Figure 3) was significantly impacted by the interaction between seawater pCO2 and temperature (statistical significance indicated by *p*-value ≤ 0.05; Table S4). Calcification rate significantly increased with pCO2 under control temperature for all three tropical species, but showed no change across pCO2 treatments under elevated temperature for *S. pistillata* or *P. damicornis*, and decreased with increasing pCO2 under elevated temperature for *S. hystrix*. Temperature had a significant negative effect on calcification rate in all tropical species (Table S4. Calcification rate of the coldwater azooxanthellate coral *L. pertusa* declined significantly with increasing pCO2 at both temperatures, but showed a significantly stronger response to pCO2 under the control temperature (Figure 3; Table S4).

**Figure 3.** The relationship between pCO2 and coral calcification rates at ambient and high temperature ((**A**) = *S. pistillata*, (**B**) = *S. hystrix*, (**C**) = *P. damicornis*, and (**D**) = *L. pertusa*). Shaded boundaries represent 95% confidence intervals. Calcification rate was significantly impacted by an interaction between pCO2 and temperature for all coral species.

#### *3.2. Calcifying Fluid Chemistry*

The pHCF of all four coral species was significantly greater than the pH of the corals' surrounding seawater (pHSW) after 30 days of exposure to ocean acidification and warming (Figure 4A–D; lmer, *S. pistillata*: *p* < 0.001; *P. damicornis*: *p* = 0.002; *S. hystrix*: *p* = 0.013; *L. pertusa*: *p* = 0.002). The pHCF of all four species declined significantly with increasing pCO2 under control temperatures (Table S5), and, also, for *S. pistillata* under elevated temperature (*S. pistillata* was the only species for which pHCF was measured at both control and elevated temperature; Figure 4). Under the elevated temperature treatment, pHCF of *S. pistillata* remained higher than pHSW, but was significantly lower than pHCF at the control temperature (Table S5). The cold-water coral *Lophelia pertusa* exhibited the steepest decline in pHCF with increasing pCO2 (Figure 5).

All four coral species increased their Δ[H+] (i.e., seawater [H+]—calcifying fluid [H+]) in response to increasing seawater pCO2 (Figure 4E–H, Table S6). The Δ[H+] of *S. pistillata* was significantly influenced by the interaction between pCO2 and temperature (Table S6). In this species, there was no difference in Δ[H+] between the two temperature treatments under control pCO2. The Δ[H+] increased with pCO2 under both temperature treatments, but this increase was significantly greater in the control temperature treatment.

**Figure 4.** The effect of seawater pCO2 on calcifying fluid pH (n = 3 individuals per treatment; panels (**A**–**D**)) and Δ[H+] (panels (**E**–**H**)) for three species of tropical corals ((**A**,**E**) = *S. pistillata*; (**B**,**F**) = *S. hystrix*, and (**C**,**G**) = *P. damicornis*) and one species of cold-water coral (*L. pertusa*; (**D**,**H**)). The impact of elevated temperature on the response of *S. pistillata* to increasing pCO2 is also shown (panels (**A**) and (**E**)). Increasing pCO2 was significantly associated with declining calcifying fluid pH and increasing Δ[H+] for all four coral species. The calcifying fluid pH of *S. pistillata* decreased significantly in response to a 3 ◦C increase in temperature, and Δ[H+] significantly responded to an interaction between increased pCO2 and increased temperature. Shaded boundaries represent 95% confidence intervals. Solid black lines represent seawater pH under control temperature; dashed black line represents seawater pH under high temperature.

**Figure 5.** Slopes of regressions calculated from linear mixed effects models investigating the impacts of pCO2 and temperature on calcification rate, calcifying fluid pH, and Δ[H+]. Significant differences were found amongst the slopes of the different coral species' calcification responses to ocean acidification and warming. No significant difference was observed amongst the slopes of the different coral species' calcifying fluid pH response to ocean acidification. The slopes of the different coral species' proton regulation response (i.e., Δ[H+]) to ocean acidification were not significantly different from each other. Vertical bars represent 95% confidence intervals.

#### *3.3. Estimated Coral Photosymbiont Index*

Both *S. pistillata* and *P. damicornis* exhibited a lower photosymbiont index (i.e., lower estimated photosymbiont abundance) under the elevated temperature treatments (Figure 6, Table S7). Photosymbiont index significantly increased in both *S. pistillata* and *P. damicornis* when pCO2 was elevated from control conditions (Figure 6, Table S7). The photosymbiont index of *S. hystrix* was significantly negatively correlated with the interaction between temperature and pCO2, whereby photosymbiont index increased significantly with increasing pCO2 under the control temperature treatment, but showed no change in response to pCO2 in the elevated temperature treatment.

**Figure 6.** The effect of seawater pCO2 and warming on the color score of the three species of tropical corals ((**A**): *S. pistillata*; (**B**): *S. hystrix*; (**C**): *P. damicornis*)). Color score is a proxy for photosymbiont abundance ('bleaching'), where 0 = bleached and 6 = healthy. Trendlines indicate significant correlations between seawater pCO2 and color score at 28 (orange) and 31 (red) ◦C. The color score of both *S. pistillata* and *P. damicornis* increased in response to increasing pCO2 (lmers, *S. pistillata*, *p* < 0.001; *P. damicornis*, *p* = 0.001) and decreased in response to a 3 ◦C temperature increase (lmers, *S. pistillata*, *p* < 0.001, *P. damicornis*, *p* = 0.001). Color score of *S. hystrix* was significantly impacted (indicating bleaching) by an interaction between pCO2 and temperature (lmer, *p* < 0.001), whereby color score increased significantly with increasing pCO2 at 28 ◦C (lmer, *p* < 0.001) but showed no significant change with increasing pCO2 at 31 ◦C (lmer, *p* = 0.056). Shaded boundaries represent 95% confidence intervals.

#### *3.4. Investigating the Role of Photosynthesis in Calcification*

The calcification rate of *S. pistillata* was best predicted by the significant interaction between photosymbiont index and temperature, whereby photosymbiont index was positively correlated with calcification rate under both temperatures, but the slope of this correlation was significantly greater under the control temperature treatment (Figure 7A, Table S8). The calcification rate of *S. hystrix* was best predicted by a model that included both photosymbiont index and temperature independently. Photosymbiont index had a positive linear relationship with calcification rate under both temperatures, but the slope of this relationship significantly decreased in the high temperature treatment (Figure 7B, Table S8). The calcification rate of *P. damicornis* was best predicted by a model that contained photosymbiont index, seawater pCO2, and temperature (Table S8). The only significant predictor of calcification rate of *P. damicornis* was the interaction between photosymbiont index, temperature, and pCO2 (Figure 7C, Table S8).

**Figure 7.** The relationship between coral color score and calcification rate of three species of tropical corals ((**A**): *S. pistillata*; (**B**): *S. hystrix*; (**C**): *P. damicornis*). Color score is a proxy for photosymbiont abundance ('bleaching'), where '0' = bleached and '6' = healthy. Trendlines indicate significant correlations between color score and calcification rate at 28 (orange) and 31 (red) ◦C. Temperature significantly impacted the relationship between color score and calcification rate of *S. pistillata*, and caused a significant decline in calcification rate of *S. hystrix*. No significant relationship existed between calcification rate and color score of *P. damicornis*, and temperature had no significant impact on this relationship. Shaded boundaries represent 95% confidence intervals.

#### **4. Discussion**

#### *4.1. Calcification Response to pCO2 and Thermal Stress*

Prior studies have shown that *S. pistillata* and *P. damicornis* both exhibit resilience in their calcification response to ocean acidification [13,51,52], while *S. hystrix* [53] tends to exhibit more negative responses. The present study, however, shows that, in the absence of thermal stress, these three common species of tropical zooxanthellate corals exhibit increased rates of calcification in response to a one-month exposure to CO2-acidified conditions. Although the disparity between the results of the past and present studies on these species could be due to differences in experimental design, such as experimental duration, the method used to measure calcification rates, temperature treatments, and/or light levels, these findings provide compelling evidence that, under certain circumstances (e.g., absence of thermal stress), some tropical zooxanthellate coral species can tolerate OA over at least one-month intervals.

Under the elevated temperature treatment (31 ◦C), the calcification rates of all three tropical coral species were reduced compared to the control temperature, but were unchanged by increasing pCO2, showing that thermal stress effectively impaired the zooxanthellate corals' calcification response to CO2-induced OA. However, it should be noted that warming can have either positive or negative effects on coral calcification rate, depending on whether the warming causes temperatures to approach or exceed, respectively, the coral's thermal optimum [19,21,54,55].

Few studies have investigated the calcification response of corals to combined ocean acidification and warming. While some of the studies show a negative response to these combined stressors [55], others contrast the results of the present study by showing no interactive effects of ocean acidification and warming [21,56,57]. The results presented here show that the impacts of ocean acidification on the calcification rates of three *Pocilloporid* coral species are exacerbated by warming. Because OA and global warming typically occur in tandem during major perturbations to the Earth's carbon cycle, both throughout Earth history [58] and as a consequence of anthropogenic CO2 emissions [59], future CO2-induced global change poses a substantial threat to these coral species.

The calcification rate of the cold-water azooxanthellate coral *L. pertusa* declined under elevated pCO2 at both temperatures. This negative calcification response to pCO2 was weaker under the elevated temperature. Notably, the direction of the pCO2-temperature interaction for the cold-water azooxanthellate species was opposite that of the three tropical species. *Lophelia pertusa* exhibited net dissolution when seawater pCO2 reached ca. 1000 ppm, although low levels of calcification have been previously observed for *L. pertusa* under similar conditions [34,35,60]. These conditions are predicted to occur in the surface open ocean by the end of the 21st century [1] and earlier in high-latitude cold-water environments [61]—suggesting that this cold-water ecosystem engineer, whose reef-systems function as nursery ground for a number of commercially important fish species [62], may be unable to form reefs beyond this century. However, the observation that increased temperature partially mitigates the impact of OA agrees with the results of longer-term studies [35], and suggests that the impacts of global change on this species will vary with temperature across latitude and depth, as this species can inhabit seawater ranging from 4 to 14 ◦C [63,64].

The differences in calcification response to ocean acidification shown here may arise from differences in the physiology and ecology of tropical vs. cold-water corals. Although tropical corals exist close to their thermal limits and are therefore vulnerable to even small amounts of warming, azooxanthellate cold-water corals can generally tolerate a wider temperature range [65]. Elevated respiration rates have been observed for *L. pertusa* in warmer temperatures [66]. Thus, an increase in temperature may boost metabolic rates to partially mitigate the impacts of OA on calcification in the high pCO2 treatments. The species of tropical corals studied here are colonial and, thus, share resources between closely packed polyps [67]. These polyps are connected by coenosarc tissue, which covers and protects the skeleton [68]. The high degree of tissue cover exhibited by tropical corals means that the skeleton is well protected from dissolution, which may explain the lack of negative calcification response (i.e., lack of net dissolution) for the tropical corals in this study. Their shared gastrovascular system allows the distribution of metabolites generated from coral respiration and zooxanthellate photosynthesis across the colony, which could yield further resilience against ocean acidification. In contrast, *L. pertusa* is a pseudocolonial species [69] that produces single polyps on the end of stalk-like branches. The coenosarc connecting these branches is often partially absent in laboratory specimens and in wild specimens during the winter, leaving regions of exposed skeleton vulnerable to dissolution [65]. This lack of connectivity between polyps and the presence of exposed skeleton may increase the vulnerability of *L. pertusa* to ocean acidification.

Alternatively, the increased solubility of CO2 in colder waters caused Ω<sup>A</sup> of the *L. pertusa* treatments to be 0.12–0.70 units lower at 9 ◦C than at 12 ◦C for a given pCO2 condition. This may have caused higher rates of skeletal dissolution in the high-pCO2 treatments maintained at the lower temperature, although the rate of dissolution of coral skeletons should be higher under the higher temperature treatments for equivalent Ω<sup>A</sup> [3]. Since the buoyant weight method [70] used here yields only a net rate of calcification, i.e., mass of new skeleton produced through gross calcification minus mass of exposed skeleton lost through gross dissolution, it is not possible to determine whether the positive impact of the interaction between pCO2 and temperature on *L. pertusa* calcification rate was driven by increased gross calcification or reduced gross dissolution (or a combination of these factors) in the high-temperature, high-pCO2 treatments. Additionally, it was not determined whether OA impacted the density of coral skeleton, which could increase the fragility of the reef framework that these corals form [71].

The responses observed in the present study are consistent with other laboratory studies showing that scleractinian corals exhibit a wide range of calcification responses to OA [13,19–24,52,72]. Some of this variability may arise from differences in experimental design, such as the amount of food provided to the corals, the levels of irradiance, and the duration of the experiment. Nevertheless, this high variability in calcification response patterns across and within species indicates that a greater understanding of the mechanisms that drive coral responses to OA is needed.

The present study was conducted over an eight-week interval, with a total of three weeks of acclimation to laboratory and experimental conditions, and should, therefore, be considered intermediate in duration. As with any experimental OA study, it is possible that the duration of exposure may influence results, as corals may function normally over short timeframes, but exhibit impaired function over longer timeframes as a result of cumulative stress and/or depletion of metabolic resources [19]. Alternatively, corals may exhibit impaired responses shortly after exposure to the treatment conditions due to shock, but acclimate to the treatment conditions over longer timeframes. Prior laboratory-based studies have shown that tropical corals exhibit variable degrees of acclimation to OA over relatively short timescales [19,24,73], whereas acclimation of *L. pertusa* has been observed over longer timescales [74]. The large inter- and intra-specific differences in coral response to OA across experiments and timeframes underscores the need for additional research into the long-term effects of OA on coral calcification.

#### *4.2. Role of Calcifying Fluid pH Regulation in Coral Response to pCO2 and Thermal Stress*

Coral calcifying fluid pH elevation has been widely cited as a mechanism for promoting calcification under conditions of ocean acidification [8,11,14]. Although pHCF declined in all four species under elevated pCO2, it always remained higher than pHSW. Notably, coral species that showed the highest degree of control over pHCF in the present experiment, and thus the shallowest slope of change in pHCF in response to changing seawater pCO2, also exhibited the greatest increase in calcification rate when pCO2 was increased, suggesting that pHCF regulation confers resilience to corals exposed to OA.

These trends are consistent with prior estimation of coral pHCF from boron isotopes [10,38,39,41,75–77], pH-sensitive fluorescent dyes [12], and pH-sensitive microsensors [9,11]. Previous studies of the effects of ocean acidification on calcification site pH show that both *S. pistillata* and *P. damicornis* elevate pHCF above seawater pH, and that this elevation increases under ocean acidification [13,52]. These results are consistent with the findings here, although the measured pHCF was considerably higher in the present study compared to previous studies.

The present study used pH-sensitive microelectrodes, whereas prior studies on both *S. pistillata* and *P. damicornis* [52], and *P. damicornis* [13] used confocal microscopy to image pH-sensitive SNARF-1 dye in coral microcolonies grown on glass slides and boron isotope systematics, respectively. Differences in pHCF could be due to differences in methods of culturing and/or pHCF-estimation, or due to genotypic differences between cultured specimens. Of the three methods used to estimate pHCF (pH-sensitive dyes, boron isotopes, and pH-sensitive microelectrodes), pH-sensitive microelectrodes typically yield the highest pHCF [11], although a side-by-side comparison of pHCF measured with pH-sensitive dye and pH microelectrodes on the same specimens yielded comparable results [14].

Measurements of pHCF using pH-sensitive microelectrodes are challenged by the difficulty in assessing the precise location of the microelectrode tip relative to the coral's calcifying fluid. This challenge was addressed in the present experiment through the creation of pHCF profiles as the pH microelectrode was withdrawn from the calcifying fluid, thus allowing characterization of the calcifying fluid pH compared to intratissue and/or gastrovascular pH (Figure 2), as was conducted in previous studies [14,78]. Additionally, the pHCF of two coral species in the present study (*S. pistillata* and *P. damicornis*) was estimated by coral skeletal δ11B to provide a side-by-side comparison of these independent approaches to measuring pHCF [41]. A significant correlation was found between pHCF

measured using these two methods, increasing confidence that the measurements obtained here represent pH of the calcifying fluid. The offset between the two approaches was attributed to the two techniques measuring pHCF over different timescales—with skeletal δ11B recording a time-averaged value of pHCF and pH microelectrodes recording a more instantaneous value of pHCF [41].

Although numerous studies have shown that OA reduces coral pHCF, few have investigated the combined impact of warming and OA on pHCF. In the present study, microelectrode measurements of the pHCF of *S. pistillata* were measured in all pCO2 and temperature treatments. The prescribed temperature increase resulted in a significant decline in pHCF for each of the three pCO2 treatments. Under heat stress, corals may receive less nourishment from their photosymbionts to support pHCF regulation due to thermally induced bleaching and/or may divert energy from the regulation of pHCF towards tissue repair.

Assuming that the coral calcifying fluid is ultimately derived from the coral's surrounding seawater [79], the extent to which a coral mitigates the impacts of OA by removing protons from its calcifying fluid can be grossly estimated (excluding the effects of buffering) from the difference between the [H+] of the calcifying fluid and the [H+] of the surrounding seawater (i.e., Δ[H+]). All four coral species exhibited increased Δ[H+] under elevated pCO2 treatments, suggesting that more energy is allocated to pHCF regulation under elevated pCO2. The Δ[H+] was lower in the high-temperature treatment in *S. pistillata*, suggesting that less energy is available to maintain elevated pHCF when the corals experience thermally induced symbiont loss—potentially due to a commensurate decrease in photosynthate translocated to the coral host.

### *4.3. Role of Photosymbionts in Coral Response to pCO2 and Thermal Stress*

In order for corals to allocate more energy toward removing protons from their calcifying fluid under OA conditions, they must divert energy from other activities and/or increase their energetic intake [80]. Zooxanthellate corals obtain energy from two sources heterotrophic feeding and photosynthate translocated from their algal symbionts. As corals are sessile suspension feeders, they are limited in the extent to which they can increase heterotrophic feeding, although they may increase feeding rates if sufficient food is available [81,82]. Alternatively, enhanced photosynthesis under OA may play an important role in driving proton elevation under OA.

Using the well-established colorimetric method of Siebeck et al. [45] to estimate the abundance of zooxanthellae via the photosymbiont index, it was shown that the populations of photosymbionts from *S. pistillata* and *P. damicornis* were significantly reduced in the high temperature treatment compared to the initial values. This is consistent with prior work showing that prolonged heat stress can lead to expulsion of zooxanthellate and tissue damage, a process termed 'bleaching' [83]. However, in this study, the photosymbiont index increased in response to increasing pCO2 under both temperature treatments, suggesting that the photosymbionts within all three *Pocilloporid* tropical corals benefited from the increased availability of dissolved inorganic carbon (DIC). The alleviation of carbon limited photosynthesis could free up energy for elevating pHCF and/or the production of carbon concentrating enzymes (e.g., carbonic anhydrase [84]), thereby aiding calcification under the control temperature [68,85,86]. The zooxanthellae's apparently positive response to increasing DIC also suggests that dissolved inorganic nutrients (DIN) were sufficient and in the correct balance (i.e., Redfield ratio) to sustain photosynthesis [87,88]. This link between enhanced photosynthesis and enhanced calcification under the elevated pCO2 and control temperature treatment is consistent with the observed correlations between photosymbiont index and calcification rate in *S. pistillata* and *S. hystrix* across all treatments.

#### *4.4. Proposed Mechanistic Framework for Zooxanthellate Coral Response to pCO2 and Thermal Stress*

We propose the following mechanistic framework to explain the zooxanthellate coral responses to OA and warming observed in this study. Under the control temperature and pCO2 conditions, zooxanthellae fix DIC as carbohydrates (photosynthate), which is then used by the coral hosts as an energy source for all physiological activities, including elevation of pHCF in support of calcification [4,8,89,90]. When OA occurs without thermal stress, high pCO2 enhances photosynthesis in the coral species investigated and increases their photosymbiont index. Under conditions of elevated pCO2, enhanced photosymbiont productivity (evidenced by increased photosymbiont index in this study), will result in a greater abundance of byproducts that may be used by the coral host to increase proton removal from the calcifying fluid (evidenced by elevated Δ[H+]). The combination of elevated pHCF and elevated DIC (due to elevated pCO2) under OA conditions may allow some species of corals to maintain an Ω<sup>A</sup> in their calcifying fluid that is comparable to or, perhaps, greater than those exhibited under non-acidified conditions [11]—hence, their observed ability to maintain constant or, in some cases, elevated rates of calcification under OA.

Although CO2-induced OA appears supportive of calcification for these three species under the control temperature, this support breaks down in the high temperature treatment. The thermal stress induced in this treatment caused a reduction in the abundance of the corals' algal symbionts (evidenced by their reduced photosymbiont index), which was accompanied by a reduction in Δ[H+] (i.e., proton removal) and calcification rate of *S. pistillata* under each of the elevated pCO2 treatments. It appears that the thermally induced reduction in photosymbiont index eliminated the benefit of enhanced photosynthesis under conditions of elevated pCO2, thereby leaving the coral with fewer resources (e.g., translocated photosynthate) for elevating pHCF. It should also be noted that thermal impairment of the enzymes used to remove protons from the coral calcifying fluid (e.g., H+/Ca2+-ATPase [9,91]) may have contributed—along with reduced photosymbiont index—to declines in pHCF and Δ[H+] observed for *S. pistillata* in response to thermal stress. Additionally, the different responses shown in the control and high temperature treatments could be due to a temperature-induced shift in the strategy used by the coral to elevate aragonite saturation state in their calcifying fluid. In natural reef systems, it has been shown that coral pHCF is most elevated in the winter months [16]. Although pHCF is still elevated in the warmer summer months, it is elevated to a lesser extent, and DIC elevation appears to be the primary means of raising Ω<sup>A</sup> [16]. These observations are consistent with and provide a potential explanation for the results of the present study.

The role of symbiotic zooxanthellae in conferring resilience to corals exposed to OA is also highlighted in the stark contrast observed between the tropical and deep-sea coral responses to OA. Cold-water corals are azooxanthellate and thus do not receive the benefits of enhanced symbiont photosynthesis under elevated pCO2. Although Δ[H+] of the coldwater species was comparable to that of the tropical species under OA, proton regulation of the calcifying fluid probably consumes a greater proportion of the total resources of the cold-water species (which acquires no resources from photosynthesis) than that of the tropical zooxanthellate species. This may leave proportionally fewer resources (compared with the tropical species) for other processes associated with calcification in the cold-water species, such as the production of organic matrices that may initiate crystal nucleation [92], which may explain the more negative calcification response to OA exhibited by *L. pertusa.*

#### *4.5. Limitations of Laboratory-Based Experiments*

The results described here were obtained in a controlled laboratory setting. In their natural reef environments, tropical corals experience fluctuations in temperature and carbonate chemistry across daily [93–95] and seasonal [96,97] cycles, whereas corals inhabiting deeper, colder waters experience more stable environments. Thus, differing degrees of prior exposure to fluctuations in pH could contribute to the differential responses of the tropical

and cold-water coral species observed here. The enhanced decline in pHCF under elevated temperature exhibited by the tropical corals could also result from a shift in strategy for elevating calcifying fluid Ω<sup>A</sup> from pHCF elevation to DIC elevation, as has been shown to occur seasonally in a natural reef system [16]. Future research is necessary to assess whether the responses observed here hold in more dynamic temperature and pH environments.

#### **5. Conclusions**

Global warming is considered to be amongst the greatest threats facing coral reefs, and OA is emerging as an equally grave threat. We show that warming has a more negative impact than OA on three species of zooxanthellate tropical corals, whereas OA has a more negative impact than warming on an azooxanthellate cold-water coral. This study also provides insight into the role of photosymbionts in corals' response to OA. Specifically, the enhancement of symbiont photosynthesis under higher-pCO2 conditions appears to mitigate the negative effects of OA on tropical zooxanthellate corals by providing resources that assist in the maintenance of elevated calcifying fluid pH in support of calcification. This resilience, however, is impaired when OA is combined with thermally induced reductions in the abundance of the coral's photosymbionts (i.e., 'bleaching'), which limits the extent to which the coral holobiont can utilize the elevated DIC via photosynthesis. These results highlight the threat that ocean warming and acidification pose for tropical and cold-water corals, especially when occurring in tandem.

**Supplementary Materials:** The following supporting information can be downloaded at: https: //www.mdpi.com/article/10.3390/jmse10081106/s1: Summary of experimental design and results (Section S1); Water quality methods and summary tables (Section S2); Buoyant weight methods (Section S3); Images of calcein dye incorporation into coral skeleton (Section S4); Method for estimating photosymbiont index (Section S5); and Model summary tables (Section S6).

**Author Contributions:** Conceptualization, J.B.R. and R.A.E.; methodology, J.B.R., R.A.E., L.P.C., D.D.B., H.W., J.V.B., G.M.S.-G. and I.W.; formal analysis, L.P.C., J.B.R. and J.G.; investigation, L.P.C., J.B.R., C.E.R., F.M.-L., I.W., J.V.B. and R.A.E. sources, J.B.R., R.A.E., D.D.B., J.B. and H.W.; data curation, L.P.C., J.B.R. and R.A.E.; writing—original draft preparation, L.P.C. and J.B.R.; writing—review and editing, L.P.C., J.B.R., R.A.E., M.G., C.E.R., J.B., J.V.B., D.D.B., G.M.S.-G. and H.W.; visualization, L.P.C., J.B.R. and J.G.; supervision, J.B.R., R.A.E., D.D.B., J.B. and H.W.; project administration, J.B.R. and H.W.; funding acquisition, J.B.R., R.A.E. and H.W. All authors have read and agreed to the published version of the manuscript.

**Funding:** J.B.R. acknowledges support from National Science Foundation grant OCE-1437371, the ZMT, and a Hanse-Wissenschaftskolleg Fellowship. R.A.E. acknowledges support from National Science Foundation grant OCE-1437166, the Pritzker Endowment to UCLA IoES, and 'Laboratoire d'Excellence' LabexMER grant ANR-10-LABX-19 co-funded by a grant from the French government under the program 'Investissements d'Avenir'.

**Institutional Review Board Statement:** Not applicable.

**Informed Consent Statement:** Not applicable.

**Data Availability Statement:** It is our intention to make data available on publication of this study.

**Acknowledgments:** We thank Artur Fink, Laurie Hoffman, and Anja Niclas (MPIMM) for assistance with construction and use of pH microelectrodes; Silvia Hardenberg, Nico Steinel, and Christian Brandt (ZMT) for assistance with coral husbandry and redesign of the acidification system at the ZMT; and Matthias Birkicht (ZMT) for assistance with water chemistry analysis.

**Conflicts of Interest:** The authors declare no conflict of interest.

#### **References**


## *Article* **Why Do Bio-Carbonates Exist?**

**Luis Pomar 1, Pamela Hallock 2,\*, Guillem Mateu-Vicens <sup>3</sup> and Juan I. Baceta <sup>4</sup>**


**Abstract:** Calcium carbonate precipitation associated with biotic activity is first recorded in Archaean rocks. The oldest putative fossils related to hydrothermal vents have been dated at ~3.77 Ga (possibly 4.29 Ga). Stromatolites, the oldest dated at 3.70 Ga, have since occurred through Earth history, despite dramatic changes in physical and chemical conditions in aquatic environments. A key question is: what advantages do photosynthesizing aquatic prokaryotes and algae gain by precipitating carbonates? We propose the Phosphate Extraction Mechanism (PEM) to explain the benefits of biomineralization in warm, oligotrophic, alkaline, euphotic environments. Carbonate precipitation enhances access to otherwise limited carbon dioxide and phosphate in such environments. This mechanism also provides an explanation for prolific production of carbonates during times of elevated atmospheric carbon dioxide at intervals in the Phanerozoic.

**Keywords:** phosphate; nutrient limitation; carbon dioxide; Archaean; Proterozoic; cyanobacteria; calcareous algae; coccolithophores

## **1. Introduction**

Carbon (C), which is the major component of fossil fuels [1], is the key ingredient in "life" [2]. Carbon is also a key ingredient in limestones (CaCO3) and other carbonate rocks [3]. The ultimate sources of carbon are the primordial components of the Earth [2]. Even at present, volcanic sources emit ~6 × <sup>10</sup><sup>8</sup> metric tons of carbon dioxide (CO2) per year [4]. Through geologic history, the dynamics of carbon chemistry have played major roles in oceanic, terrestrial and even subsurface-crustal processes [2]. Human activities are now contributing roughly 40 billion tons of CO2 annually into the Earth's atmosphere [5], with climatic consequences that are becoming more apparent every year.

The systematic measure of CO2 concentrations in the atmosphere was started by C. David Keeling of the Scripps Institution of Oceanography in March 1958 at a NOAA (National Oceanic and Atmospheric Administration) facility at Mauna Loa Observatory, Hawai'i, USA [6]. NOAA initiated its own CO2 measurement in May 1974 and, since then, NOAA and Scripps have run the measurements in parallel [7]. Prior to Keeling's research, CO2 measurements were inconsistent, but the Keeling's methods revealed a pattern of seasonal oscillations of CO2, with peaks in May and lows in November, with the averaged values of successive years progressively increasing (Figure 1). Keeling envisaged the yearly cycles to reflect the vegetation cycles that prevail across the northern hemisphere while the increase over time was thought to be caused by human activities, especially the burning of fossil fuels [8].

**Citation:** Pomar, L.; Hallock, P.; Mateu-Vicens, G.; Baceta, J.I. Why Do Bio-Carbonates Exist?. *J. Mar. Sci. Eng.* **2022**, *10*, 1648. https://doi.org/ 10.3390/jmse10111648

Academic Editors: Hildegard Westphal, Justin Ries and Steve Doo

Received: 27 September 2022 Accepted: 21 October 2022 Published: 3 November 2022

**Publisher's Note:** MDPI stays neutral with regard to jurisdictional claims in published maps and institutional affiliations.

**Copyright:** © 2022 by the authors. Licensee MDPI, Basel, Switzerland. This article is an open access article distributed under the terms and conditions of the Creative Commons Attribution (CC BY) license (https:// creativecommons.org/licenses/by/ 4.0/).

**Figure 1.** CO2 concentrations in the atmosphere [7], started by C. David Keeling of the Scripps Institution of Oceanography in March 1958 [8]. (Adapted with permission from Scripps CO2 program, Scripps Institution of Oceanography at UC San Diego, CA, USA.).

By the 1960 s, greenhouse gas emissions and their link to global climate change became a serious concern, with both scientists and the public considering CO2 to be an invisible pollutant. Attention and concern for greenhouse gases and their role in climate change intensified, and in 1988 the United Nations created the Intergovernmental Panel on Climate Change (IPCC) [9]. Concern for ocean acidification, *climate change's equally evil twin*, emerged somewhat later. Although scientists had been tracking ocean pH for more than 30 years, biological studies emerged in the 1990s [10] and have accelerated since the introduction of the term "ocean acidification" in 2003 [11].

Burning of fossil fuels, combined with widespread changes in land use, has resulted in rapidly increasing concentrations of atmospheric CO2 that are causing the decline in the pH of surface seawater [10–14]. Will progressive warming and acidification cause mass extinctions and evolutionary turnover of marine biotas, as occurred at previous events that caused major perturbations of ocean chemistry [15–17]?

This question reveals a major paradox in the geologic record. Through the Phanerozoic, extended Greenhouse World times of higher CO2 (Silurian–Devonian and Jurassic– Cretaceous) were characterized by prolific accumulations of biogenic carbonates, especially those associated with cyanobacterial and algal calcification (Figure 2).

**Figure 2.** Carbonate precipitation nodes, genetic macroevolution rates, and estimates of oxygen and CO2 in the atmosphere. (**A**): Supercontinents and crustal growth; precratonic, accretionary orogens and supercontinents assembly (redrawn from multiple sources). (**B**): Carbonate precipitation modes

through Earth's history (multiple sources). [1]: Oldest fossils [18]. [2]: Isua oldest stromatolites [19]. (**C**): pH evolution from [20]; [**a**]: nominal model, in which the median Archean surface temperature is slightly higher than modern surface temperatures. Archean land fraction was anywhere between 10% and 75% of modern land fraction; [**b**]: no Archean land endmember scenario; [**c**]: model with assumed 100 ppm Proterozoic methane and 1% Archean methane levels. (**D**): Maximum and minimum estimates of atmospheric partial oxygen pressure (after [21]). (**E**): Estimates of atmospheric CO2 concentration; [3] from [22], upper and lower boundaries reflect average surface temperature for an ice-free (20 ◦C) and ice covered (5◦) Earth; [4] Acritarch isotopic composition [23]; [5] C-isotope reservoir modeling [24]; [6] Phanerozoic GEOCARB III models [25]; [7] Royer compilation [26]; [8] from [27]; [9] paleosoil mass balances [28]; [10] from [29]; [11] Picoplanktonic whiting and partial sheath calcification commenced 1400–1300 Ma ago (33 CO2 PAL), and cyanobacteria CCMs were induced when pCO2 = 10 PAL [30]. PAL: present atmospheric level; GOE: the Great Oxygenation Event; NOE: Neoproterozoic oxygenation event. (This figure is an original compilation and interpretation of data by L. Pomar, based upon numerous sources).

For example, during the Cretaceous, micritic and algal (e.g., coccoliths) production was prolific. The emergence of rudists as major metazoan producers of skeletal carbonates was notable, and the "elevator" nature of many lineges has been postulated to have been the response to rapid inundation by carbonate muds. In the Cenozoic, the Paleogene was a time of transition. Coccolithophores and planktic foraminifers emerged as major producers of pelagic carbonate sediments through the Mesozoic. Following the Cretaceous-Paleogene extinction, new lineages of coccolithophores and planktic foraminifers diversified, while calcareous macroalgae and larger benthic foraminifers were notable producers of neritic carbonates. Coral reefs emerged as major carbonate factories as Icehouse World conditions developed in the Oligocene and became predominant in the Neogene, a conundrum noted more than 45 years ago [31].

Another paradox associated with changes in ocean chemistry over the course of Earth history relates to another essential element for life, which is commonly the most limiting nutrient required for essential processes including photosynthesis, growth and reproduction. Bioavailable phosphorus (P), typically occurring as phosphate (PO4 <sup>3</sup>−), provides the backbone of nucleotides (DNA and RNA) and phospholipid membranes of cells, and is used in many other essential functions including energy storage and transfer associated with adenosine triphosphate (ATP), and for the synthesis of proteins and enzymes. Kempe and Degens [32] postulated that abundant dissolved PO4 <sup>3</sup><sup>−</sup> in the Archaean seas helped to foster the evolution of life (see also [33,34]).

#### **2. The Hypothesis**

The goal of our paper is to explore the hypothesis that decline in both CO2 and PO4 <sup>3</sup>−, and the history of changes in modes of calcification through Earth history, are tied to the close and complex relationships between biogenic carbonates and phosphate availability. The **Phosphate Extraction Mechanism** (**PEM**) provides an explanation for how calcifying microbial and algal biota can thrive in oligotrophic conditions. During daylight, calcification can be coupled to photosynthesis. With energy from active ion transport, protons can be split from bicarbonate, providing a carbonate ion (CO3 <sup>2</sup>−) for calcification and a CO2 molecule for photosynthesis. Phosphate is adsorbed during calcium carbonate (CaCO3) precipitation. At night, respiration releases CO2 and can promote partial dissolution of diurnally precipitated carbonates. Phosphate that is adsorbed onto precipitating CaCO3 during daylight can be desorbed at night, making it available for uptake by the calcifying cyanobacteria or algae.

#### **3. The Hadean–Archaean**

Conditions on the Hadean Earth were clearly very different from present, and even from those of the Proterozoic. One hypothesis is that the early Earth was hot, following the moon-forming impact, and cooled to the point where liquid water was present after about 10 million years [35,36]. Whether the Earth's surface continued to be thermophilic well into the Archaean is still debated, though ∼4.3 Ga rocks near Hudson Bay are suspected to have formed under warm, greenhouse conditions [37].

Over the past 4.5 Ga, heat-, gravity-, and tectonically driven processes have concentrated Iron (Fe) and Nickel (Ni) in the Earth's core. The dominant elements in the Earth's crust are Oxygen (O), Silicon (Si), Aluminum (Al), Iron (Fe), Calcium (Ca), Magnesium (Mg), Sodium (Na), and Potassium (K). In the mantle, Mg is nearly an order of magnitude more prevalent than in the crust, where proportions especially of Si and Al are higher. Basaltic, komatiitic (ultramafic) and mantle ejecta are potent CO2 sinks (all are rich in Mg2+ and Fe2+, with somewhat less Ca2+) and their subduction may have drawn down atmospheric CO2 [36,38]. Alkalic rocks in India indicate carbonate subduction occurred by 4.26 Ga [36].

Through the Archaean, the combination of high atmospheric CO2 concentrations and the prevalence of Mg2+, Fe2+, and Ca2+ resulted in formation of massive abiogenic dolomites and limestones, indicating extreme oversaturation of waters [39] (Figure 2B). The earliest hints of life date back ~3.8 billion years (early–mid Archaean, Figure 2A,B), and by ~3.7 Ga, some metacarbonate rocks, found in the Isua supercrustal belt in southwest Greenland, contain 1–4-cm-high isolated and aggregated stromatolites [19]. Through the Proterozoic, massive abiogenic carbonates declined in prevalence relative to biogenic carbonates and are unknown in modern marine environments, even in those that are strongly oversaturated with respect to CaCO3. In a water body containing abundant Ca2+ and dissolved inorganic carbon, modest increases in saturation, whether by warming, reduced hydrostatic pressure, evaporation, or by biogenic processes such as photosynthetic uptake of CO2 or microbial sulfate reduction, can trigger CaCO3 or CaMg(CO3)2 precipitation.

So the questions we pose are these: In post Archaean oceans, did biogenic processes become so ubiquitous that they became the drivers and controllers of CaCO3 precipitation, resulting in the overwhelming predominance of biogenic carbonate production? Was the sequence of carbonate factories through geologic history, from predominantly abiogenic precipitation in the Archaean, to predominance of biologically induced geochemical precipitation in the Proterozoic, to the addition of biologically controlled calcification during the Phanerozoic (Figure 2B), consistent with and driven by evolutionary processes involving luminosity of the Sun, the stratification of the Earth, and the emergence of life, with their combined effects on ocean chemistry (Figure 2C)?

The luminosity of the Sun has increased ~30% during the past 4.5 Ga [40]. The composition of the Earth's atmosphere has changed (Figure 2D,E), with substantial decline in greenhouse gases, including CO2 (Figure 2E) and methane (CH4). For example, CO2 has been estimated to have been ~100 times higher in the Hadean than modern levels and ~10 times higher in the Cambrian than present [36]. As noted above, ultramafic ejecta are potent CO2 sinks and their subduction may have drawn down atmospheric CO2 in the Hadean [38].

#### **4. Pertinent Chemical Processes**

Before delving into biological processes, we briefly examine some chemical factors associated with CO2 in aqueous environments (Figure 3). Two measures associated with but not restricted to CO2 are pH and alkalinity. Both are important in water chemistry, but their interactions can be confusing when considering calcification (29,38,41). The concentration of hydrogen ions (H+), indicating how acidic or basic a substance is, is measured as pH. Alkalinity indicates the ability of a solution to neutralize acids (i.e., buffering capacity). Alkalinity consists of ions that can incorporate H+ (protons) into their structure, limiting the availability of those protons that would otherwise lower pH. Carbonate alkalinity measures the concentrations of CO3 <sup>2</sup><sup>−</sup> and HCO3 − ions, which are typically present in the highest concentrations in natural waters. Total Alkalinity reflects primarily CO3 <sup>2</sup><sup>−</sup> and HCO3 − ions, but also includes PO4 <sup>3</sup>−, borate, orthosilicate, sulfides, and organic acids. Shallow water bodies can be strongly influenced by rising temperature and evaporation, concentrating the alkaline ions and promoting the precipitation of carbonate minerals, assuming that appropriate cations are available (e.g., Ca2+, Mg2+). Moreover, any process that takes up protons can also promote such precipitation.

**Figure 3.** In an aqueous solution, carbonate, bicarbonate, carbon dioxide, and carbonic acid exist together in dynamic equilibrium, e.g., [10,20]. When CO2 is absorbed by seawater, a series of chemical reactions increase the concentration of H+ and cause the seawater acidity to increase. When calcium carbonate (e.g., limestone) reacts with acidic free hydrogen (H+) ions in seawater, the solid calcium carbonate dissolves, releasing free calcium (Ca2+) ions and free bicarbonate (HCO3 −) ions. (This figure is an original interpretation by L. Pomar based on numerous sources).

As elements that are relatively abundant in both crustal and mantle rocks, Mg2+, Fe2+, and Ca2+ ions are soluble in aquatic environments. Thus, the combination of high carbonate alkalinity and abundant reactive cations, especially in warm waters, provided prolific sources of ions to produce abiogenic carbonates, providing critical storage of atmospheric CO2 on geological time scales [41].

#### **5. Pertinent Biological Processes**

All living organisms carry traces of the histories of their ancestors within their genetic makeup and, in the case of the Eukarya, literally within their cells. The diversity of metabolic pathways in the prokaryotic Archaea and Eubacteria far surpasses the comparatively limited pathways found in the Eukarya, which are metabolically limited to those of their Proteoarchaeota and α-proteobacterial symbiotic predecessors [42].

Early microbial forms evolved in aquatic environments lacking free oxygen (O2). Autotrophic processes in aquatic systems require: (a) an energy source, (b) source of dissolved inorganic carbon (DIC), (c) a source of protons, and (d) nutrients (fixed N, P, Fe, and trace elements) necessary for cell growth and reproduction [43]. Energy sources include oxidation of inorganic compounds (e.g., H2S, CH4) by chemoautotrophs and sunlight by photoautotrophs [44]. The primary sources of DIC are CO2 or CH4. In photoautotrophy, proton donors can be H2, Fe2+ or H2S, which do not release O2 (anoxygenic photosynthesis), or H2O that releases O2 as a byproduct (oxygenic photosynthesis).

RuBisCO is the enzyme that catalyzes the first major step of carbon fixation in the Calvin cycle, the process of photosynthesis. As the most abundant protein on Earth, RuBisCO is found in all three domains of life (Eubacteria, Archaea and Eukarya), and fixes more than 90% of the inorganic carbon that is converted into biomass [45]. However, this enzyme evolved in an O2-free atmosphere, prior to the Great Oxygenation Event, and does not efficiently discriminate between CO2 and O2. Enzyme activity and specificity are reciprocally linked: faster RuBisCO has a higher error rate and more specific RuBisCO has a lower catalytic rate [45].

#### **6. CCMs: CO2 Concentrating Mechanisms**

It may seem paradoxical that RuBisCO, this ubiquitous and essential enzyme, has not become more efficient or been replaced by more efficient enzymes. However, CCMs (CO2 concentrating mechanisms) evolved instead by adaptation of active transport processes and compartmentalization of accumulated HCO3 − [46]. The CCMs in cyanobacteria and microalgae enhance photosynthetic carbon fixation by energy-driven uptake of HCO3 − and its conversion to CO2. The minimum requirements for a cyanobacterial CCM are one energy-driven active transporter that accumulates HCO3 − in the cytosol, and an anionpermeable carboxysome containing intracellular carbonic anhydrase [47]. Central to the functioning of the cyanobacterial CCM is the carboxisome (Figure 4), a protein microcompartment within the cell that contains RuBisCo and a carbonic anhydrase. The latter converts HCO3 − into CO2 within the carboxysome, concentrating CO2 up to 1000-fold around the active site of RuBisCO [46], thus overcoming the inefficiency of RuBisCO [24] and its inability to distinguish O2 from CO2.

**Figure 4.** Central to the cyanobacterial CCM (CO2 Concentrating Mechanism) is the carboxisome, a cellular micro-compartment containing RuBisCo and a carbonic anhydrase (CA) that, by converting HCO3 − into CO2, concentrates CO2 as much as 1000-fold around the active site of RuBisCO [46]. (This figure is an original interpretation by L. Pomar based on several sources).

For eukaryotic algae with CCMs, the presence of pyrenoids is often associated with the occurrence of a CCM, although there are exceptions [47]. Eukaryotic algal CCMs likely originated independently in different clades of algae, as indicated by the diversity of inorganic C transporters and of carbonic anhydrases, as well as the lack of a strict correlation between occurrence of a CCM and the presence of a pyrenoid [47].

#### **7. Phosphorous**

Another essential element for life forms that has fluctuated downward in bioavailability through Earth history is phosphorous (P). Ranking 11th among the elements in the Earth's crust, P occurs primarily as phosphate (PO4 <sup>3</sup>−) in apatite. Although found in most igneous and metamorphic rocks, mafic rocks commonly contain at least an order of magnitude more apatite than most felsic rocks (granites and rhyolites) [29]. Thus, as Feand Mg-rich materials have been concentrated in the Earth's mantle and core, so too has P. By the late Archean, continental-scale granitic cratons had developed on the Earth's crust, contributing to the decline in PO4 <sup>3</sup><sup>−</sup> in aquatic environments [32].

Phosphorus is one of the eleven macro-biogenic elements that make up living organisms [C, O, H, N, S, P, Na, K, Ca, Mg, Cl, Fe]. Bioavailable PO4 <sup>3</sup><sup>−</sup> is essential for all forms of life, although it has received less attention than C, N, and Fe in the evolution of autotrophy in relation to environmental changes during the last 4 Ga [48–50]. As noted in the Introduction, PO4 <sup>3</sup><sup>−</sup> is essential to cellular structure and processes, including nucleotides (DNA and RNA), phospholipid membranes, energy transfer and storage, and for synthesis of proteins and enzymes [32]. However, PO4 <sup>3</sup><sup>−</sup> became one of the most limiting nutrients for primary productivity in Phanerozoic marine ecosystems due to its low solubility in oxygenated waters and because, unlike bioavailable dissolved N (DIN), PO4 <sup>3</sup><sup>−</sup> cannot be fixed from the atmosphere [34,48,51].

Microbial uptake of PO4 <sup>3</sup><sup>−</sup> takes place through at least two kinetically distinct processes. A "low affinity component" operates continuously, apparently driven by protonmotive force. A "high affinity component" requires energy from ATP and operates when internal PO4 <sup>3</sup><sup>−</sup> pools are depleted. Was the decline in available PO4 <sup>3</sup><sup>−</sup> in Proterozoic oceans related to changes in major precipitation modes of CaCO3?

A key factor in the bioavailability of PO4 <sup>3</sup><sup>−</sup> in the Archean and early Proterozoic oceans was the lack of free oxygen [42,52]. By the early Proterozoic, cyanobacteria were producing free oxygen, triggering the Great Oxidation Event (~2.4–2.1 Ga) (Figure 2). With the emergence of O2 into the atmosphere, terrestrial weathering dramatically changed, resulting in oxidation of sulfides ubiquitous in continental rocks. A major result was the delivery of substantial amounts of sulfates to the oceans, and those dissolved sulfates provided the oxidation potential for sulphate-reducing microbes, which are obligate anaerobes. The result was an oxygenated mixed layer/photic zone to perhaps 200 m depth, but the vast ocean depths remained anoxic and sulphidic (i.e., euxinic) through most of the Proterozoic [42,52].

The bioavailability of PO4 <sup>3</sup><sup>−</sup> declined in surface waters, based on the relative insolubility of PO4 <sup>3</sup><sup>−</sup> in oxygenated waters compared to its much greater solubility in euxinic waters [51,52]. Shallow waters exposed to sunlight became depleted in PO4 <sup>3</sup><sup>−</sup> as a consequence of reduced solubility in terrestrial-runoff waters and in oxygenated surface waters, and were further diminished by biological uptake by photosynthetic cyanobacteria and microalgae, and by geochemical uptake into ferric oxyhydroxides. At the same time, the much greater volumes of subsurface, aphotic waters, which were euxinic and therefore in which PO4 <sup>3</sup><sup>−</sup> was highly soluble, became major repositories. In that respect, like today, hydrodynamic processes such as mixing and upwelling must have been critical mechanisms for supplying PO4 <sup>3</sup><sup>−</sup> to photosynthesizing microorganisms in the surface waters [51].

By the late Neoproterozoic, sufficient photosynthetic oxygen production had occurred to oxidize most oceanic waters. Two additional contributors to the increase in atmospheric O2 and dissolved O2 throughout the oceans were primary production by eukaryotic algae and increased solubility of O2 in cold oceanic waters associated with Cryogenian glaciation. The latter also would have increased the rates of ocean turnover and delivery to surface waters of dissolved PO4 <sup>3</sup><sup>−</sup> from subsurface waters (Figure 5). Shen et al. [53] postulated that the oxygenation of the oceans allowed diverse animal lineages to increase in size.

On scales of major plate-tectonics processes, massive basalt flows and tectonic uplift increase delivery of weathering-mobilized PO4 <sup>3</sup><sup>−</sup> to the oceans [48,49]. On intermediate time scales, changes in rates of deep-ocean circulation influence rates of delivery of DIN and PO4 <sup>3</sup><sup>−</sup> to surface waters, where they promote primary productivity [51,54]. In some cases, regional changes in ocean circulation have brought PO4 <sup>3</sup><sup>−</sup> rich subsurface waters into oxygenated conditions, creating major phosphatization events that have resulted in economically valuable phosphate ore deposits such as the prolific Lutetian-Messinian phosphorites of Morocco [55]. Most PO4 <sup>3</sup><sup>−</sup> in the ocean is cycling within the biotic system, and is returned to seawater by microbial activity. Only a small fraction of the phosphorus is buried in sediments, as reduced conditions within the sediments promotes dissolution of PO4 <sup>3</sup><sup>−</sup> into sediment pore waters and its return to ocean cycling [54].

**Figure 5.** Evolution of Phosphorous (P) concentration through the late Neoproterozoic and Cambrian and its relationship with the emergence of major groups of organisms. Notice that the first weakly calcified metazoans (red bar) appeared during a minimum of P at the end of the Ediacaran. Modern lineages of metazoans with hard skeletons appeared by mid-Cambrian during another minimum in P concentration. (This figure is an original interpretation by L. Pomar based on numerous sources).

#### **8. Calcification as a Sink and a Source of Both CO2 and PO4 3**−

Precipitation of carbonate minerals is a long-term sink for CO2, but in the short term, can serve as a source (e.g., [41] and references therein). Calcification typically involves two bicarbonate ions, incorporating one CO3 <sup>2</sup><sup>−</sup> into the CaCO3 mineral structure and releasing the other into the environment as CO2 [56]. In aquatic environments, photosynthetic uptake of CO2 or HCO3 − reduces availability of CO2, increases pH and can lead to oversaturation with respect to calcite. Dissolved PO4 <sup>3</sup><sup>−</sup> is adsorbed onto and co-precipitates with calcite, with incorporation of some of the surface PO4 <sup>3</sup><sup>−</sup> into the bulk structure as crystal growth proceeds (Figure 6). However, higher concentrations of dissolved phosphate inhibit the growth of calcite crystals and can stop growth completely [57].

Numerous studies have documented removal of PO4 <sup>3</sup><sup>−</sup> associated with CaCO3 precipitation, for example during dense blooms of some cyanobacteria. Guldbransen and Cremer [58] experimentally demonstrated co-precipitation of CaCO3 and PO4 <sup>3</sup>−, and Kitano et al. [59] found that stirring increases the amount of PO4 <sup>3</sup><sup>−</sup> that co-precipitates. Millero et al. [60] demonstrated that PO4 <sup>3</sup><sup>−</sup> uptake on CaCO3 minerals is a multistep process. Depending upon water chemistry and temperature, adsorption or desorption can occur rapidly, in minutes to a few hours, with much slower processes lasting more

than one week. Up to 80% of the adsorbed PO4 <sup>3</sup><sup>−</sup> can be released from CaCO3 over one day. The amount of PO4 <sup>3</sup><sup>−</sup> left on the CaCO3 is close to equilibrium adsorption [60]. Thus, the adsorption of PO4 <sup>3</sup><sup>−</sup> during CaCO3 precipitation can be both a sink and a source of phosphates (Figure 7).

**Figure 6.** Simplified diagram of the adsorption process (adapted from [58–60]). Atoms, molecules or ions adhere to a surface by weak residual forces (e.g., van der Waals, electrostatic). It can be multilayer or monolayer. (This figure is an original interpretation by L. Pomar based on several sources).

**Figure 7.** Co-precipitation of phosphate and calcium carbonate; note that phosphate is more readily adsorbed onto and released from aragonite than onto/from calcite ((This figure is an original interpretation by L. Pomar based on several sources, including [58–60]).

#### **9. Prokaryotic Organo-Sedimentary Systems**

Stromatolites, the fossil evidence of calcified microbial mats, can be viewed as the first biotic process to produce a hard body [61]. By the Late Archaean, ooids—the second biotic process to produce a hard body—appeared when photosynthetic O2 production by cyanobacteria was occurring. What might have been the advantage for these benthic microbial systems to produce coated-mineral structures? Such coatings would have restricted light penetration, as well as diffusion of DIC and nutrients required for photosynthesis, growth and reproduction. Furthermore, what could be the advantage of inducing carbonate-mud precipitates (whitings) by neritic photosynthetic prokaryotes (picoplankton) and eukaryotes (phytoplankton)? This third microbially mediated organo-sedimentary system to produce carbonate precipitates emerged by the Mesoproterozoic, following the Great Oxygenation Event (GOE) (Figure 2).

Cyanobacteria have remarkably few nutritional requirements. They use light as source of energy, H2O or H2S as electron donors and CO2 (or HCO3 −) as the source of inorganic carbon [62]. Additionally, many species can fix atmospheric N2 to NH4 +, which makes them independent of DIN sources such as nitrate, ammonium or organic nitrogen. Hence, the most critical nutrients for cyanobacteria are PO4 <sup>3</sup>−, which can be stored intracellularly as polyphosphate, and, to a lesser extent, Fe2+. Both are limited by their reactivity with O2. Moreover, cyanobacteria can minimize their PO4 <sup>3</sup><sup>−</sup> requirements through the synthesis of sulfolipids [63], allowing them to colonize low-nutrient environments, as demonstrated by their ubiquity in both terrestrial and aquatic environments.

Thus, precipitation of CaCO3 during the day augments CO2 availability for photosynthesis [64], while PO4 <sup>3</sup><sup>−</sup> adsorption sequesters that scarce but essential nutrient. At night, desorption of PO4 <sup>3</sup><sup>−</sup> makes it available for uptake by the cells. In microbial mats (Figure 8) in darkness, anoxic conditions resulting from continued sulfate reduction, aerobic respiration and sulfide oxidation likely contribute to the dissolution of CaCO3 and increase the potential for PO4 <sup>3</sup><sup>−</sup> uptake by the microbial assemblage.

**Figure 8.** Oxygen and sulfide concentration (ppm) in (**A**) and pH in (**B**) variation during a diel cycle in the upper 12 mm of modern marine stromatolites in the Exuma Cays, Bahamas. (**C**): Light intensity and temperature at the surface of the stromatolite. (Adapted with permission from Ref. [65]. Copyright ©2002, American Chemical Society.).

Giant ooids, associated with cyanobacterial mats and stromatolites, are common in Late Archaean rocks. They have been described in the 2.64 Ga old Ghaap Group in South Africa [66] and in the 2.72 Ga old Pilbara group in Australia [67], becoming abundant and characteristic of many Neoproterozoic carbonates [68]. Microbially mediated organomineralization processes are likely involved in the origin of ooids. They can either be biologically induced (e.g., by-products of metabolic activities that increase environmental alkalinity and trigger carbonate precipitation), or biologically influenced (e.g., microbial extracellular polymeric substances can serve as templates for carbonate mineralization [69]). Again, the advantage for the microbial consortium to precipitate CaCO3 coatings around particles was likely in PO4 <sup>3</sup><sup>−</sup> adsorption, associated with the precipitation of metastable amorphous carbonate [70,71] (Figure 9).

**Figure 9.** Conceptual multiphase model for the oolite cortex accretion. Abbreviations: ACC (amorphous CaCO3) and EPS (extracellular polymeric substances). (This figure is an interpretation by L. Pomar based on cited sources [70,71]).

A significant source of lime-mud production in modern oligotrophic shallow marine and lacustrine environments also is linked to microbial calcification [72]. Whitings are the drifting milky clouds that are commonly observed in warm, shallow seas. Robbins et al. [73] estimated that whitings production might account for more than 40% of the bank-top- and peri-platform Holocene muds on the west side of Great Bahama Bank. The production of whitings is linked to blooms of cyanobacteria and planktic green algae in oligotrophic aquatic environments via CO2 uptake [74]. Photosynthetic carbon fixation during blooms induces an increase in extracellular pH, favoring crystal nucleation [72]. What would it be the advantage for photosynthetic pico- and microplankton to induce carbonate-mud precipitation in the water? Blooming means that CO2 and nutrient acquisition must keep pace with the increase of organic matter required for cell division. With the diurnal cycle of rapid adsorption and desorption of PO4 <sup>3</sup>−, the milky cloud can efficiently capture this scarce nutrient, which would be of particular advantage in extremely oligotrophic conditions.

#### **10. Biocalcification in Photosynthetic Eukaryotes**

The basic eukaryotic cell is descended from an anaerobic Proteoarchaeota that engulfed purple, non-sulfur α-proteobacteria that could utilize oxygen but were not obligate aerobes [42,44]. All algae have that basic eukaryotic cell and all algal plastids are believed to have originated from cyanobacterial endosymbionts [42,75]. Molecular phylogenetic studies have demonstrated that the morphological diversity of the algae results from their polyphyletic origins within the Eukarya [75,76]. An interesting question then, is: Why do some marine algae within different phylogenetic groups invest energy into producing CaCO3 skeletal structures, whereas most do not?

Tropical macroalgae cope with three strong selective pressures: limited access to CO2/HCO3 − due to warm water and competition among photosynthetic organisms, limited access to PO4 <sup>3</sup>−, and herbivory. Calcification is clearly an anti-herbivory defense. Furthermore, there are striking differences in percent tissue PO4 <sup>3</sup><sup>−</sup> between calcified and fleshy algae, indicating different nutrient-acquisition strategies. The effects of nutrient enrichment with nitrogen (DIN) and PO4 <sup>3</sup><sup>−</sup> on productivity and calcification of fleshy and calcareous algae differ. DIN and PO4 <sup>3</sup><sup>−</sup> frequently enhance productivity of fleshy algae, but do not increase calcification rates of calcareous species [77].

*Halimeda*, a codiacean macroalga common in warm, shallow-marine environments, can add up to one segment per day, per branch (Figure 10). However, the algae require nutrient uptake for growth and reproduction in warm waters where PO4 <sup>3</sup><sup>−</sup> levels are often below detection limits. Calcification occurs during daytime within an organic matrix in the outer utricle walls of one-day-old segments. Photosynthetic removal of CO2 by diffusion across a membrane and enzymatic anhydrase activity induces aragonite precipitation in the utricle space [78,79]. At night, respiration elevates CO2 and decreases CO3 <sup>2</sup><sup>−</sup> saturation, which decreases seawater pH in the utricle space. The aragonite needles partially dissolve, break and recrystallize into micro-anhedral crystals (<1 μm). Such partial dissolution at night would facilitate desorption of PO4 <sup>3</sup><sup>−</sup> from the skeleton.

Coralline red algae occur as thin crusts on hard substrata that can produce free-living rhodoliths, or as thalli with articulated branches. Most precipitate high-Mg calcite rather than aragonite. Calcification dynamics (Figure 11) depend on nutrient concentrations; PO4 <sup>3</sup><sup>−</sup> inhibits the crystalline lattice and hampers calcite precipitation [80]. Enrichment of PO4 <sup>3</sup>−, along with other nutrients, favors the growth of macroalgae and phytoplankton, which reduces water transparency and impedes coralline algal growth. Calcification associated with photosynthesis exhibits day-night cyclicity that is also recorded in diel pH fluctuations. Abiotic re-precipitation of CaCO3 can occur in crustose-coralline red algae during nighttime [81] and pH drop at the boundary layer in darkness has been documented [82]. Thus, partial dissolution occurring at night would provide favorable conditions to remobilize adsorbed PO4 <sup>3</sup><sup>−</sup> from the calcitic skeleton and facilitate uptake by the algal cells.

**Figure 10.** Calcification stages of *Halimeda* [78,79]. (**A**): Growth of short skeletal needles on the utricle wall; (**B**): recrystallization of short needles to micron-sized anhedral crystals. (**C**): secondary calcification by long, dense, euhedral, skeletal-aragonite needles, (**D**): the primary inter-utricular space in the rim of the segment is filled with micron-sized anhedral crystals. (This figure is an original interpretation by L. Pomar based on cited sources [78,79]).

**Figure 11.** Skeletal formation in red algae. Process steps: (1) Primary cell wall is forming. Hemicellulose microfibrils act as nucleating substrate for micro-granules. (2) Cellulose microfibrils extrude from plasma membrane and leaks through broken primary cell wall. (3) Interfilament grains initiate attached to the external part of the primary cell wall and grow in the leaked cellulose microfibril. (4) Secondary cell wall thickens, extruded cellulose microfibrils stop leaking and mineralize as Mg-calcite. (5) Plasmalemma is pushed inward as cellulose microfibrils continue to extrude and mineralize. Secondary cell wall fully formed. (This figure is an original interpretation by L. Pomar based on sources including [80]).

Coccolithophores are photosynthetic protists that produce small calcitic disks (coccoliths). These oval-shaped plates of CaCO3 consist of double discs, composed of radial

arrays of minute, elaborately-shaped crystal units. They form delicate crystalline lace or open web-like patterns in which the rate of the CaCO3 crystal surface/volume is very high. Originating by the Late Triassic (Figure 2), they proliferated during the Cretaceous, and, together with planktic foraminifera, allowed widespread carbonate production in the open ocean [83]. Coccolithophores can produce as much as two coccoliths per hour [84]. The variability of coccolith shapes indicates they serve diverse functions. Young [85] noted features that defy simple physical functions, particularly complex mesh structures and the variety of elaborate forms, and suggested that coccoliths might be adaptations for nutrient uptake.

Coccolithophores thrive in waters with minimal PO4 <sup>3</sup>−, that is, insufficient to promote growth of other phytoplankton. Coccolith formation is less adversely affected by nutrient deficiency than is cell division and growth [86]. The nutrient requirements for the organic cellular components of a coccolithophorid cell are similar to those for non-calcifying phytoplankton. However, forming coccoliths requires minimal nutrient cost, as coccolith production continues even when cell division ceases because of nutrient limitation [87,88].

In modern *E. huxleyi*, coccoliths nucleate and grow in a Golgi-derived coccolith vesicle, from where they migrate to the cell surface (Figure 12). Within the coccolith vesicle, polysaccharides are thought to regulate nucleation and subsequent growth of coccoliths [89]. Sviben et al. [90] identified a reservoir compartment, distinct from the coccolith vesicle, with high concentrations of Ca2+ and PO4 <sup>3</sup>−. Only Ca2+ is delivered to the coccolith vesicle [91], whereas PO4 <sup>3</sup><sup>−</sup> is diverted to the cell for growth and reproduction.

**Figure 12.** Calcification processes in coccolithophorids. [**A**] Coccolith formation occurs within the Golgi-derived coccolith vacuole (CV). [**B**] Intracellular CaCO3 precipitation from bicarbonate (HCO3 <sup>−</sup>) releases equimolar H<sup>+</sup> that must be rapidly removed from the CV to maintain pH for CaCO3 precipitation. [**C**] Some H+ may be utilized by photosynthesis. [**D**] Though not fully understood, it is hypothesized that Ca2+ is recruited through Ca-channels and V-type ATPase pumps. [**E**] A reservoir compartment (RC) concentrates Ca++ and PO4 <sup>3</sup>−. [**F**] The coccolith-associated polysaccharide may drive the Ca2+ from the coccosphere into the RC, from where only Ca2+ is transported to the coccolith vesicle (CV) and PO4 <sup>3</sup><sup>−</sup> is diverted to the cell for algal growth and reproduction. Transfer of Ca2+ from the RC into the CV uses Ca2+ channels and transporters. (This figure is an original interpretation by L. Pomar based on [88–91] and other sources).

#### **11. CO2 and Carbonate Depositional History**

In aqueous solutions, CO3 <sup>2</sup>−, HCO3 −, CO2, and H2CO3 co-exist in dynamic equilibria (Figure 3). When CO2 reacts with water (H2O), it complexes to H2CO3, which dissociates to H<sup>+</sup> and HCO3 <sup>−</sup>, thereby increasing the concentration of H<sup>+</sup> and cause the seawater acidity to increase (i.e., the pH to decline). When CaCO3 (e.g., limestone) reacts with free hydrogen (H+) ions in seawater, the solid CaCO3 dissolves, forming free calcium (Ca++) ions and free bicarbonate (HCO3 −) ions. As noted previously, the concentrations of ions of carbonate, bicarbonate, phosphate, borate, orthosilicate, sulfides, and organic acids constitute alkalinity in aquatic environments, especially HCO3 −. The consumption of protons results in decreased H+ activity. Thus, CaCO3 acts to neutralize or buffer the solution by consuming H+.

In the modern world, burning of fossil fuels, combined with widespread changes in land use, have resulted in rapidly increasing concentrations of atmospheric CO2 that are causing the decline in the pH of surface seawater (e.g., [6–11]). Will the progressive warming and acidification cause mass extinctions and evolutionary turnover of marine biotas, as has been documented by paleoclimatic evidence [15–17]?

From this question emerges another major "carbonate paradox". Contrary to predictions of increased carbonate dissolution in response to increasing CO2, extended high CO2 Greenhouse World periods during the Phanerozoic (Silurian–Devonian and Jurassic– Cretaceous) were characterized by thick and extensive accumulations of fine-grained carbonates associated with cyanobacterial and algal calcification, commonly baffled by skeletal carbonates (e.g., stromatoporids, corals, rudists) (Figure 2) ([92] and references therein). In the Cenozoic, coral reefs emerged as major carbonate factories as Icehouse World conditions developed and became predominant in the Oligocene through the Holocene, a conundrum noted by Frost many years ago [31].

The solution to this paradox lies in two parts. First, there is the "rate" factor. Mass extincton events in the Phanerozoic record have long be recognized by carbonate depositional hiatuses that indicate times of global carbonate dissolution [93]. Very rapid increases in atmospheric CO2 (Figure 3) or other atmospheric compounds that unite with water to form acids (e.g., 2SO2 + H2O→H2SO4 + 2H+), whether caused by massive extrusion of flood basalts, bolide impact, methane-hydrate release, extensive burning of fossil fuels, or some combination of events, can indeed trigger carbonate dissolution resulting in hiatuses of tens of thousands of years (Paleocene–Eocene) or hundreds of thousand years (Cretaceous– Paleocene) [94] and references therein) or millions of years (Permian–Triassic) [17].

However, eventually terrestrial weathering and dissolution under high atmospheric CO2 over time increased alkalinity and restored the carbonate factory. The massive accumulations of Silurian–Devonian and Jurassic–Cretaceous carbonates occurred during times of relatively rapid sea-floor spreading, which elevated both CO2 and Ca2+ concentrations, triggered warming temperatures, sea-level rise, and expansive areas of relatively shallow basins and shelves. Reduced land areas limited freshwater input of dissolved nutrients and warm waters reduced rates of deep-ocean circulation. During Greenhouse-world conditions, the trophic resource continuum expanded (compared to Icehouse conditions), with limited extremely rich regions of upwelling and expansive regions of extreme oligotrophy, the latter fostering extensive production of carbonates [95]. Elevated evaporation rates, relatively shallow seas, and high alkalinity favored calcifying cyanobacterial and algal taxa that could sequester scarce PO4 <sup>3</sup><sup>−</sup> while producing excess organic matter upon which calcifying animals (e.g., rudists) could also thrive [92]. The name Cretaceous was actually derived from the Latin *creta*, meaning chalk [96].

During daylight hours, in warm, shallow waters, abundant microalgae take up available DIN and PO4 <sup>3</sup>−, while high rates of photosynthesis deplete immediate access to CO2/HCO3 −. Calcification associated with photosynthesis alleviates both deficiencies. Two HCO3 − provides one CO2 for photosynthesis and a CO3 <sup>2</sup><sup>−</sup> for calcification, somewhat alleviating local CO2 depletion. At the same time, PO4 <sup>3</sup><sup>−</sup> adsorption occurs during daytime calcification and is partially desorbed at night, becoming available for uptake by the

cyanobacteria or algae. In addition, some dissolution of the CaCO3 releases HCO3 −, and along with community respiration, results in early morning maximum concentrations of HCO3 −. With the onset of light, photosynthesis makes energy available to CCMs, providing the primary producers with access to and storage of CO2, which becomes increasingly unavailable over the course of the day.

Finally, why did extensive coral reefs only become prevalent with the onset of the Icehouse World climates [31]? Again, the answer is likely climatic and geochemical influences on biological processes [92,97–100]. High-latitude cooling, compression of tropical habitats, and increasing temperature gradients between high and low latitudes and between surface and deeper waters [97] were major factors. Coral communities flourished circumtropically beginning in the late Oligocene [31]. By the early Miocene, reefs and associated biota expanded latitudinal distributions by more than 10◦ north and south [98]. The expansion in reef-building capacity corresponded to increasing Mg/Ca ratios in seawater [99] and falling atmospheric CO2 concentrations [100], both of which promoted aragonite precipitation in warm, tropical waters. The majority of extant Symbiodiniacea lineages diversified since Middle Miocene [101]. Global cooling, including much cooler subsurface waters, likely benefited the coral-Symbiodiniacea symbioses, which are sensitive to photo-oxidative stress under elevated temperatures [102]. The substantial increase in rates of circulation of deep, cold waters resulted in more "intermediate" conditions of oceanic oligotrophy [95]. Pomar and Hallock [92] further postulated that Neogene cooling supported the co-evolution and expansion of zooxanthellate corals and coralline algae into shallow, high-energy waters, where their carbonate production potential was highest. In such environments, both calcification and PO4 <sup>3</sup><sup>−</sup> extraction are also potentially optimized.

#### **12. Conclusions**

Photosynthesis typically involves daytime CO2 uptake and O2 release predominating over respiration; at night, respiration consumes O2 and releases CO2. The result is strong diel variation in both pH and dissolved O2 concentrations in aquatic environments. In warm-water environments in which alkalinity is relatively high and PO4 <sup>3</sup><sup>−</sup> concentrations are minimal, these variations have the potential to promote precipitation of CaCO3 and associated adsorption of PO4 <sup>3</sup><sup>−</sup> in daylight when photosynthesis is active. At night, lower pH and oxygen availability can promote some CaCO3 dissolution and PO4 <sup>3</sup><sup>−</sup> desorption. The combination of calcification and PO4 <sup>3</sup><sup>−</sup> extraction occurs in photosynthesizing cyanobacteria, as well as in a diverse array of calcifying nanophytoplankton and calcareous macroalgae, indicating that this process allows both prokaryotic and eukaryotic photosynthetic organisms to thrive in warm, alkaline, oligotrophic waters. Concentrations of both atmospheric CO2 and surface-ocean PO4 <sup>3</sup>−, which are essential for biological productivity, have declined by at least 1–2 orders of magnitude over Earth history. **The Phosphate Extraction Mechanism**, associated with photosynthetically induced calcification, has played a major role in the production and accumulation of carbonates throughout much of the Proterozoic and the Phanerozoic. Recognition of this relationship helps resolve the apparent paradox that periods in the Phanerozoic when atmospheric CO2 levels were considerably higher than present (as well as higher than predicted Anthropocene concentrations) were times of massive accumulation of carbonates, predominantly produced by photosynthesizing cyanobacteria and calcifying algae. This synthesis also stresses the importance of rates of change, as times of rapid increase in atmospheric CO2 concentrations were associated with mass-extinction events characterized by global carbonate-depositional hiatuses, requiring hundreds of thousands or even millions of years for surface waters of the oceans to regain sufficient alkalinity to sustain accumulation of massive carbonates.

**Author Contributions:** Conceptualization: L.P., P.H., G.M.-V. and J.I.B.; Methodology: L.P., P.H., G.M.-V. and J.I.B.; Investigation: L.P., P.H., G.M.-V. and J.I.B.; Writing: Original Draft Preparation, L.P. and P.H.; Writing: Review and Editing, P.H. and G.M.-V. All authors have read and agreed to the published version of the manuscript.

**Funding:** JIB acknowledges funding from the Basque Government to the Research Group IT1602-22. L.P., P.H. and G.M.-V. participation did not involved external funding aside from their academic institutions.

**Institutional Review Board Statement:** Not applicable.

**Informed Consent Statement:** Not applicable.

**Data Availability Statement:** Not applicable.

**Acknowledgments:** Gabriel Moyà (UIB) for the seminal discussions on this topic, and to the volume editor, H. Westphal, for her patience and suggestions.

**Conflicts of Interest:** The authors declare no conflict of interest.

#### **References**


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