Why Do Bio-Carbonates Exist?

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.

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 × 10⁸ 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].

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 [10,11,12,13,14]. 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 [15,16,17]?

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).

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 (PO₄³⁻), 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 PO₄³⁻ 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 PO₄³⁻, 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 (CO₃²⁻) 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 CO₃²⁻ and HCO₃⁻ ions, which are typically present in the highest concentrations in natural waters. Total Alkalinity reflects primarily CO₃²⁻ and HCO₃⁻ ions, but also includes PO₄³⁻, 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.

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 anion-permeable carboxysome containing intracellular carbonic anhydrase [47]. Central to the functioning of the cyanobacterial CCM is the carboxisome (Figure 4), a protein micro-compartment 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.

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 (PO₄³⁻) 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 Fe- and 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 PO₄³⁻ 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 PO₄³⁻ 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 [48,49,50]. As noted in the Introduction, PO₄³⁻ 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, PO₄³⁻ 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), PO₄³⁻ cannot be fixed from the atmosphere [34,48,51].

Microbial uptake of PO₄³⁻ takes place through at least two kinetically distinct processes. A “low affinity component” operates continuously, apparently driven by proton-motive force. A “high affinity component” requires energy from ATP and operates when internal PO₄³⁻ pools are depleted. Was the decline in available PO₄³⁻ in Proterozoic oceans related to changes in major precipitation modes of CaCO₃?

A key factor in the bioavailability of PO₄³⁻ 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 PO₄³⁻ declined in surface waters, based on the relative insolubility of PO₄³⁻ in oxygenated waters compared to its much greater solubility in euxinic waters [51,52]. Shallow waters exposed to sunlight became depleted in PO₄³⁻ 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 PO₄³⁻ 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 PO₄³⁻ 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 PO₄³⁻ 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 PO₄³⁻ to the oceans [48,49]. On intermediate time scales, changes in rates of deep-ocean circulation influence rates of delivery of DIN and PO₄³⁻ to surface waters, where they promote primary productivity [51,54]. In some cases, regional changes in ocean circulation have brought PO₄³⁻ 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 PO₄³⁻ 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 PO₄³⁻ into sediment pore waters and its return to ocean cycling [54].

8. Calcification as a Sink and a Source of Both CO₂ and PO₄³⁻

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 CO₃²⁻ 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 PO₄³⁻ is adsorbed onto and co-precipitates with calcite, with incorporation of some of the surface PO₄³⁻ 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 PO₄³⁻ associated with CaCO₃ precipitation, for example during dense blooms of some cyanobacteria. Guldbransen and Cremer [58] experimentally demonstrated co-precipitation of CaCO₃ and PO₄³⁻, and Kitano et al. [59] found that stirring increases the amount of PO₄³⁻ that co-precipitates. Millero et al. [60] demonstrated that PO₄³⁻ 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 PO₄³⁻ can be released from CaCO₃ over one day. The amount of PO₄³⁻ left on the CaCO₃ is close to equilibrium adsorption [60]. Thus, the adsorption of PO₄³⁻ during CaCO₃ precipitation can be both a sink and a source of phosphates (Figure 7).

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 PO₄³⁻, 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 PO₄³⁻ 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 PO₄³⁻ adsorption sequesters that scarce but essential nutrient. At night, desorption of PO₄³⁻ 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 PO₄³⁻ uptake by the microbial assemblage.

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 organo-mineralization 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 PO₄³⁻ adsorption, associated with the precipitation of metastable amorphous carbonate [70,71] (Figure 9).

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 PO₄³⁻, 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 PO₄³⁻, and herbivory. Calcification is clearly an anti-herbivory defense. Furthermore, there are striking differences in percent tissue PO₄³⁻ between calcified and fleshy algae, indicating different nutrient-acquisition strategies. The effects of nutrient enrichment with nitrogen (DIN) and PO₄³⁻ on productivity and calcification of fleshy and calcareous algae differ. DIN and PO₄³⁻ 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 PO₄³⁻ 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 CO₃²⁻ 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 PO₄³⁻ 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; PO₄³⁻ inhibits the crystalline lattice and hampers calcite precipitation [80]. Enrichment of PO₄³⁻, 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 PO₄³⁻ from the calcitic skeleton and facilitate uptake by the algal cells.

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 PO₄³⁻, 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 PO₄³⁻. Only Ca²⁺ is delivered to the coccolith vesicle [91], whereas PO₄³⁻ is diverted to the cell for growth and reproduction.

11. CO₂ and Carbonate Depositional History

In aqueous solutions, CO₃²⁻, 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⁺ and HCO₃⁻, thereby increasing the concentration of H⁺ 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., [6,7,8,9,10,11]). Will the progressive warming and acidification cause mass extinctions and evolutionary turnover of marine biotas, as has been documented by paleoclimatic evidence [15,16,17]?

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 PO₄³⁻ 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 PO₄³⁻, 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 CO₃²⁻ for calcification, somewhat alleviating local CO₂ depletion. At the same time, PO₄³⁻ 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 [92,97,98,99,100]. 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 PO₄³⁻ 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 PO₄³⁻ concentrations are minimal, these variations have the potential to promote precipitation of CaCO₃ and associated adsorption of PO₄³⁻ in daylight when photosynthesis is active. At night, lower pH and oxygen availability can promote some CaCO₃ dissolution and PO₄³⁻ desorption. The combination of calcification and PO₄³⁻ 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 PO₄³⁻, 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.