Kawahata et al. Progress in Earth and Planetary Science
https://doi.org/10.1186/s40645-018-0239-9(2019) 6:5

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Progress in Earth andPlanetary Science

Open Access

Perspective on the response of marine
calcifiers to global warming and ocean
acidification—Behavior of corals andforaminifera in a high CO2 world “hot house”
Hodaka Kawahata1,2*, Kazuhiko Fujita3, Akira Iguchi2, Mayuri Inoue4, Shinya Iwasaki5, Azumi Kuroyanagi6,
Ayumi Maeda1, Takuya Manaka7, Kazuyoshi Moriya8, Haruka Takagi1, Takashi Toyofuku5, Toshihiro Yoshimura5
and Atsushi Suzuki2

Abstract
The CO2 concentration of air has increased over the last two centuries and recently surpassed 400 ppm. Carbon
cycle models project CO2 concentrations of 720 to 1000 ppm for the IPCC intermediate scenario (RCP 6.0), resulting
in an increase in global mean temperature of ~ 2.6 °C and a decrease in seawater pH of ~ 0.3. Together, global
warming and ocean acidification are often referred to as the “evil twins” of climate change, potentially inducing
severe threats in the near future. In this paper, our discussion is focused on the response of two major calcifiers,
foraminifera and corals, which contribute much to the global carbonate burial rate. Photosymbiosis is regarded as
an adaptive ecology for living in warm and oligotrophic oceans, especially for reef-building corals and larger reefdwelling benthic foraminifera. As a consequence of global warming, bleaching may be a global threat to algal
symbiont-bearing marine calcifying organisms under conditions of high temperature and light intensity. If CO2 isdissolved in seawater, the partial pressure of CO2 in seawater (pCO2) and dissolved inorganic carbon (DIC) increases
while pH and the saturation state of carbonate minerals decreases without any change in total alkalinity. Generally,marine calcifying organisms show decreases in calcification rates in response to acidified seawater. However, theresponse often differs depending on situations, species, and life-cycle stage. Some benthic foraminifera showed apositive response to low pH conditions. The Acropora digitifera coral calcification of adult branches was not reduced
markedly at higher pCO2 conditions, although calcification tended to decrease versus pCO2 in both aposymbiotic
and symbiotic polyps. New analytical technologies help identify important constraints on calcification processes.
Based upon Ca isotopes, the transport path of Ca2+ and the degree of its activity would predominantly control the
carbonate precipitation rate. Visualization of the extracellular pH distribution shows that proton pumping produces
the high internal pH and large internal–external pH gap in association with foraminiferal calcification. From theperspective of a long-term change in the Earth’s surface environment, foraminifera seem to be more adaptive androbust than corals in coping with ocean warming and acidification but it is necessary to further understand themechanisms underlying variations in sensitivity to heat stress and acidified seawater for future prediction. Since CO2
is more soluble in lower temperature seawater, ocean acidification is more critical in the polar and high-latitude(Continued on next page)* Correspondence: kawahata@aori.u-tokyo.ac.jp
1
Atmosphere and Ocean Research Institute, The University of Tokyo, 5-1-5
Kashiwanoha, Kashiwa, Chiba 277-8564, Japan
2
Geological Survey of Japan, National Institute of Advanced Industrial
Science and Technology (AIST), Higashi 1-1-1, Tsukuba, Ibaraki 305-8567,
Japan
Full list of author information is available at the end of the article© The Author(s). 2019 Open Access This article is distributed under the terms of the Creative Commons Attribution 4.0
International License (http://creativecommons.org/licenses/by/4.0/), which permits unrestricted use, distribution, andreproduction in any medium, provided you give appropriate credit to the original author(s) and the source, provide a link to
the Creative Commons license, and indicate if changes were made.

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

Page 2 of 37(Continued from previous page)

regions. Additionally, older deep-water has enhanced acidity owing to the addition of CO2 from the degradation oforganic matter via a synergistic effect with high pressure. With current ocean acidification, pH and the saturation
state of carbonate minerals are decreasing without any change in total alkalinity. However, in the Earth’s history, it
is well known that alkalinity has fluctuated significantly. Therefore, it is necessary to quantitatively reconstructalkalinity, which is another key factor determining the saturation state of carbonate minerals. The rapid release ofanthropogenic CO2 (in the present day and at the Paleocene/Eocene boundary) induces severe ocean acidification,
whereas in the Cretaceous, slow environmental change, even at high levels of pCO2, could raise alkalinity, thereby
neutralizing ocean acidification.
Keywords: Partial pressure of CO2, Global warming, Ocean acidification, Coral, Foraminifera, Bleaching, Calcite,
Aragonite, Saturation state, Organic matter, Alkalinity, Carbon cycle

Introduction
Human beings are now dramatically changing our planet
and altering the environment by modifying the global carbon cycle through anthropogenic carbon dioxide (CO2)emissions from the burning of fossil fuels. The CO2 concentration of air was around 280 ppm in 1750 AD (the beginning of Industrial Revolution), but it has increased over thelast two centuries and is currently above 400 ppm. Carbon
dioxide is one of the primary greenhouse gases in Earth’s atmosphere, together with water vapor, methane, nitrous
oxide, and others. The Intergovernmental Panel on Climate
Change (IPCC) intermediate RCP 6.0 scenarios (IPCC
2013b) predicts that the CO2 concentration in air will increase to 720–1000 ppm by 2100 AD and that the global
mean temperature will rise by 2.6 °C. Comparison of CO2
concentrations in the atmosphere with fossil-fuel burning
data indicates that about 60% of fossil-fuel emissions remainin the air. The remainder is mainly absorbed by the ocean,which may provide some relief from severe warming.
Atmospheric carbon reservoirs are controlled by the formation/dissolution of both carbonate and organic matter(OM) in the ocean, which changes the chemical equilibrium of the marine CO2 system by removing/adding alkalinity in association with the release/absorbance of CO2 in
the atmosphere (Fig. 1) (e.g., Kleypas et al. 1999; Kleypas
et al. 2006; Orr et al. 2005). In open ocean, biogenic carbonate production is positively correlated with biogenic
OM production, although the carbon ratios of OM to carbonate generally increase as OM fluxes increase (based
upon the results from sediment trap experiments in the
Pacific Kawahata 2002). Planktic foraminifera, which livein the ocean surface layers from tropical to polar regions,secrete low-Mg calcite shells and contribute 32–80% of
the total deep-marine calcite budget in the global carbonate cycle (Schiebel 2002). The comparable carbonate contributor in the open ocean is coccolithophore. This is aunicellular, eukaryotic phytoplankton (alga) and one of themore abundant primary producers in the ocean, especially in tropical and subtropical oceans influenced by

El Nino/Southern Oscillation phenomena (e.g., Beaufort
and Grelaud 2017).
In oligotrophic environments such as coral reefs and subtropical gyres in the Pacific, the symbiotic relationship between the host and its symbiotic algae presents three majoradvantages: energy from photosynthesis, an enhancementof calcification, and uptake of host metabolites. Bleaching isdefined as (i) damage to the symbiotic relationships between hosts and algal symbionts and (ii) photosynthetic
pigment loss (Iglesias-Prieto et al. 1992). Elevated seawatertemperature brings negative effects on marine calcifiers; inparticular, coral bleaching has occurred repeatedly since
1979 AD (e.g., Bunkley-Williams and Williams 1990) in
tropical regions worldwide (Brown 1997; HoeghGuldberg et al. 2007). Mass bleaching events such as thosein 1979–1980 AD, 1982–1983 AD, and 1986–1988 AD
provide important opportunities for investigating the influence of bleaching during the early stages of this research area. Recently, bleaching has been found in benthic
foraminifera in coral reefs (Schmidt et al. 2011). Coral
reefs are among the ecosystems most vulnerable to environmental stresses associated with future climate changes
caused by human activities (IPCC 2017).
By the end of twenty-first century, the surface waters of
the ocean could be more acidic, resulting in an undersaturation condition with respect to carbonate minerals in someareas of high latitude. Ocean acidification (OA) will induce
severe stress in calcifiers, some of which are sensitive to environmental changes at global and local levels. Additionally,
Oschlies et al. (2008) pointed out that anthropogenic CO2
emissions will extend oxygen-minimum zones, a 50% increase in the suboxic (defined as dissolved O2 < 5 mmol m−3)
water volume by the end of this century, although a serious
ocean anoxic event, which occurred in the Cretaceous, will
not occur in the normal open ocean. Drivers such as OA,
higher seawater temperature, or reduced dissolved oxygen
concentration may overlap to amplify ecosystem impacts.
Therefore, understanding the effect of environmental
parameters (i.e., temperature, salinity, pH, and dissolved

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

aPage 3 of 37bFig. 1 a Schematic diagram of modern carbon cycle for both global warming and ocean acidification, which are often referred to as the “eviltwins” of climate change. The increase in atmospheric CO2 concentration caused by human activities is predicted to promote global warming by~ 2.5 °C and an ocean pH reduction of ~ 0.3 by 2100 AD. Coral reefs are among the ecosystems most vulnerable to environmental stresses(bleaching and/or ocean acidification). An aragonite saturation state of 3.3 is suggested to be the environmental limit for the existence of tropicalcoral reefs where calcification exceeds the dissolution of calcium carbonate (Kleypas et al. 2006). The culture experiments reported general
negative effects on survival, calcification, growth, and reproduction despite significant variation among taxonomic groups in marine environments. For
example, calcification tended to decrease versus pCO2 in both aposymbiotic and symbiotic polyps of Acropora digitifera, although its coral calcificationof adult branches was not reduced so much under higher pCO2 conditions. In contrast to the majority of benthic foraminifers dwelling in coral reefs,
Calcarina gaudichaudii showed a positive response to increases in pCO2. Since CO2 is more soluble in lower temperature seawater, ocean acidification
is more critical in the polar and high-latitude regions. Surface water in the Southern Ocean will be undersaturated with respect to aragonite by2050 AD (Orr et al. 2005). More severe dissolution of carbonate will occur in deep-sea due to the increase in solubility at higher pressure, which will
make the CCD shallow. Most likely, deep-sea benthic foraminifera will be forced to change from calcareous to the agglutinated forms reported at the
Paleocene/Eocene boundary. b Schematic diagram of the carbon cycle in coral reefs. The equations involve the following important processes: (1)photosynthesis (CO2 invasion, resulting in slowing acidification), (2) respiration/degradation (CO2 evasion, promoting acidification), (3) calcification (CO2
evasion), and (4) dissolution of carbonate (CO2 invasion, one of the neutralizing processes). Processes (1) and (2) are associated with organic carbon
metabolism, whereas (3) and (4) are associated with inorganic carbon processes. Photosynthesis triggers the production of OH−, commonly resulting
in the deposition of carbonate. In contrast, the reaction of carbonate with CO2 works as a sink of CO2. This reaction is often promoted in deep-sea,
which results in the depletion of carbonate on abyssal plains in the Pacificoxygen concentration) on representative calcifiers such ascorals and foraminifera is essential for the verification ofcarbonate cycles and future environments. In the global
carbonate budget, the carbonate production rates in theocean and the tropical shelves are estimated to be ~ 1.0 Pg
C year−1 and ~ 0.17 Pg C year−1, respectively. However, itsburial rate to sediments in the ocean is ~ 0.15 Pg C year−1
via 85% dissolution in water column/surface sediments;
this is comparable to that in the tropical shelves(~ 0.13 Pg C year− 1). Foraminifera, coccolithophores, andpteropods accounted for 43%, 22%, and 34%, respectively,in the ocean (personal communication with Prof. Tim
Rixen). Of these, pteropods often produce aragonite shells,
which are more liable to dissolution. Our discussion is focused on two major calcifiers, foraminifera and corals, although OA may impact other groups of marine organisms(e.g., Iglesias-Rodriguez et al. 2002; Catubig et al. 1998;
Gangsto et al. 2008; Tambutté et al. 2011).
In this review, initially, we describe the general featuresof the marine carbonate system and biotic and abiotic

carbonate minerals. In the next section, we summarize theresponse of corals and foraminifera to OA. Photosymbiosis is regarded as an adaptive ecology for living in warmand oligotrophic oceans and also enhances calcification of
the host’s skeleton (e.g., Stanley Jr. 2006; Moya et al.
2008). Therefore, it is believed to facilitate the prosperityof calcifying organisms in such environments. After discussing the role of symbiosis, we report the latest resultson coral and foraminiferal bleaching including culture
experiments, and then discuss the responses of
symbiont-bearing calcifiers in a warm and/or high CO2
world. New chemical and isotope analyses of biominerals,
together with visualization techniques for extracellular pH
distribution, provide valuable information regarding biomineralization including calcification. Furthermore, thelatest results on carbonate dissolution using X-ray
micro-computed tomography will be presented. Finally,
we discuss results on the response of calcifiers to different
environments such as extreme warmth in the Cretaceous,the middle Eocene climatic optimum (MECO) event, and(2019) 6:5

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perturbations in the global carbon cycle at the Paleocene/
Eocene boundary. This discussion includes the interaction
between terrestrial and marine environments to better
understand the key factors inducing OA in the past andthe near future.

Foraminifera, corals, haptophytes, mollusks, and sea
urchins produce skeleton structures of carbonate, mainlyin the mineral form of aragonite and high and low magnesium (Mg) calcites. Of these, foraminifera, corals, andhaptophytes are the most important organisms forming
large amounts of carbonate rock, calcareous sediment,
and coral reef; thus hosting one of the richest and mostdiverse ecosystems of our planet (Kleypas et al. 1999).
Both carbonate formation and OM production play important roles in controlling the amount of CO2 in theoceans and atmosphere. Generally speaking, OM production and degradation work as a sink and a source ofCO2, respectively. Conversely, carbonate formation in
the sea via the combination of calcium ions (Ca2+) andHCO3− liberates CO2, whereas its dissolution works as asink of CO2 (Fig. 2). Additionally, the chemical weatheringof terrestrial rocks plays an important role in the long-term
global carbon cycle (Sarmiento and Gruber 2006). The
weathering of both carbonate and silicate rocks consumes
atmospheric CO2 and generates bicarbonate (HCO3−),
which is transported to the ocean and released back to theatmosphere in the form of CO2 through oceanic carbonate
mineralization processes (Eq. 3 in Fig. 1):

Kawahata et al. Progress in Earth and Planetary Science

Background to ocean acidification
When CO2 is dissolved in seawater, chemical reactions reduce seawater pH, carbonate ion concentration ([CO32−]),and the saturation states of biogenic carbonate minerals(Fig. 1):CO2ðaqÞ þ H2 O⇄H2 CO3 ⇄HCO3 ‐‐þ Hþ ⇄CO3 2 þ 2 Hþ

ð1ÞThese chemical reactions are termed “ocean acidification (OA).” The seawater pH of surface ocean waters
was 8.17 at 1750 AD but this has decreased to 8.06
in the present day and is expected to reach to 7.93
(pCO2 = 560 μatm) or 7.71 (pCO2 = 1000 μatm) in thenear future (Gattuso et al. 1998). OA is known as“climate change’s equally evil twin”, in reference tothe significant and harmful consequence of enhanced
CO2 in the atmosphere. As the pH scale is logarithmic, a change of 1.0 pH unit represents approximately a 30% increase in acidity.

CaSiO3 þ 2CO2 þ H2 O→Ca2þ þ 2 HCO3 ‐ þ SiO2

ð2Þ

CaCO3 þ CO2 þ H2 O→Ca2þ þ 2 HCO3 −ð3ÞFig. 2 Total alkalinity and dissolved inorganic carbon (DIC) diagrams showing contours of pH (a), partial pressure of CO2 (b), and saturation statesof aragonite (Ωarag; c) and calcite (Ωcal; d). Chemistry of surface water under modern conditions (solid red circle) at St. 96 (5°00’N, 144°55.7’E;
data are available from http://pacifica.pices.jp/table/) will be modified in 2300 AD (open circle)

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Silicate and carbonate weathering in this sequence of processes consumes the difference in the molar ratio of CO2.
Over long timescales (104 to 106 years), it is suggested that
only silicate weathering contributes to a net reduction of atmospheric CO2 concentrations, which are closely linked topast climate conditions (Berner 1992; Brady 1991).

shown in the contours in the DIC and TA plane because these parameters are functions of DIC and TA(Figs. 2 and 3). The four parameters, DIC, TA, pH, andpCO2, are all measurable parameters in the marine CO2
system. More importantly, OA is expressed as a horizontal
vector in the Deffeyes’ diagram (DIC–TA plot), as shownin Fig. 2. OA is the result of increased oceanic CO2 uptake
by continuing anthropogenic emissions. CO2 dissolution
in seawater increases DIC but causes no changes in TA(Table 1).
The conditions of under- and super-saturation of seawater with respect to calcium carbonate (CaCO3) are represented by calcium carbonate solubility (Ω), which isdefined as the ratio of the product of the solutes over the
product of the solutes at saturation (solubility product):Kawahata et al. Progress in Earth and Planetary Science

Review
Basics of ocean acidification

OA is in essence a CO2-related phenomenon in the ocean.
Therefore, its basic physicochemical characteristics can bewell illustrated by chemical equilibrium modeling of theCO2 system in seawater (Figs. 2, 3, 4, and 5, Table 1).
General features of the carbonate system

Dissolved inorganic carbon (DIC = [CO2] + [HCO3−] +[CO32−]) and total alkalinity (TA = [HCO3−] + 2[CO32−] +[B(OH)4−] + [OH−] + 2[PO43−] + [HPO42−] + [SiO(OH)3−]– [H+] – [HSO4−]) are the most important parameters in
the aquatic CO2 system because they are essential for representation of the laws of conservation of mass and electricity. A cross plot of DIC and TA is known as the
Deffeyes’ diagram (Deffeyes 1965). pH and pCO2 can be



 
 

Ω ¼ CO3 2‐ Ca2þ = CO3 2‐ sat Ca2þ sat
¼ CO3 2‐ Ca2þ =Ksp

ð4Þ

where [CO32−]sat and [Ca2+]sat are, respectively, the concentrations of carbonate and dissolved Ca2+ in equilibriumwith carbonate mineral (CaCO3). [CO32−] and [Ca2+] are
the observed CO32− and Ca2+ concentrations. [CO32−] is

a

b

c

d

e

f

Fig. 3 Upper panels: changes in pCO2 in seawater and the corresponding saturation states of calcite and aragonite (Ωcal and Ωarag) as seawater
pH decreases under three different climate schemes: a polar and high-latitude regions (2 °C), b temperate (15 °C), and c tropical (28 °C). Lower
panels: changes in [CO32−] with the progression of ocean acidification at three different temperature settings of 2 °C (d), 15 °C (e), and 28 °C (f).
The corresponding DIC and [HCO3−] are shown in each graph. The chemical equilibrium calculation was conducted using CO2calc (Robbins et al. 2010),
with the following conditions: total alkalinity of 2200 μmol kg− 1, salinity of 34, and pressure of 10 dbar

Kawahata et al. Progress in Earth and Planetary Scienceab(2019) 6:5

cPage 6 of 37deFig. 4 Depth profiles of temperature and salinity (a), total alkalinity and DIC (b), pH and total hydrogen concentration scale (c), and saturation
states for calcite (Ωcal; d), and aragonite (Ωarag; e) at St. 96 (5° 00′ N, 144° 55.7′ E; water depth 4102 m; sampling date April 22, 1995) in the tropicalwestern Pacific during the NOPACCS NH95-1 cruise (data were downloaded from http://pacifica.pices.jp/table/). The compositions of the surface
water include a temperature of 29.4 °C, salinity of 33.859, DIC of 1870 μmol kg− 1, and total alkalinity of 2225 μmol kg−1. The chemical equilibrium
calculation was conducted using CO2calc (Robbins et al. 2010) to elucidate temperature and pressure influences on changes in the depth profilesof Ωcal and Ωarag. Note that changes in CO2 system chemistry are responsible for rapid decreases in Ωcal and Ωarag below the surface layer, whereaspressure influence becomes dominant at greater depth. Ksp is much dependent on the pressure: Ksp for calcite = (1.95–0.0176 × T)10− 2 × S × 10((0.693–
0.00645 × T)10–4 × z)
; Ksp for aragonite = (3.06–0.0217 × T)10− 2 × S × 10((0.650–0066 × T)10–4 × z)

closely related to the acidity of seawater. Ksp is the solubility product, which is dependent on temperature, salinity,pressure, pH, and nutrient levels. The term Ω is also
known as the carbonate saturation state and is calculated
for both calcite and aragonite (Ωcal and Ωarag, Robbins
et al. 2010). When Ω > 1, seawater is supersaturated withrespect to carbonate minerals; conversely, when Ω < 1,
seawater is undersaturated. Because [Ca2+] in the ocean
has a long residence time (106 years), it can be assumed tobe constant on a time scale shorter than 104 years. [Ca2+]
is assumed to be proportional to the salinity level. Therefore, the carbonate saturation state (Ω) is mainly determined by [CO32−] and Ksp (Table 1).
The saturation states of calcite and aragonite (Ωcal andΩarag) can be calculated and shown in the DIC–TA plot
(Figs. 4 and 5). OC modifies the CO2 system, and its change
is represented by the horizontal vector in the DIC-TA plot.
It always results in a decrease in Ω (Table 1).
While OA may progress ubiquitously, the corresponding changes in pH and pCO2 vary locally. The most evident parameter to influence the pace of OA progression
is temperature. By using the parameters for the carbonate system in the equatorial Pacific as representative,
Fig. 3 summarizes the temperature effects on marine

CO2 system parameters, including the saturation statesof carbonate minerals (Table 1). Although pCO2 is sensitive to temperature (~ 4% increase per 1 °C), this cannot
be seen in the upper panels of Fig. 3. In contrast, Ωcaland Ωarag show large variations according to thetemperature settings (Fig. 3a–c). Ωcal of the 2 °C condition becomes lower than the critical value of 1 (carbonate dissolution) when pCO2 reaches around 750 μatm or
pH 7.7 (Fig. 3c). In contrast, Ωcal for the 28 °C condition
may not reach the critical value before pCO2 reaches
around 2300 μatm or pH 7.3 (Fig. 3c), clearly indicatingthat OA is more critical in the polar and high-latitude
regions compared with temperate and tropical regions.
This trend become easily understandable in light of thetemperature dependency of [CO32−], as shown in thelower panels in Fig. 3d–f.
Another important factor influencing the saturation
states of seawater with respect to carbonate minerals ispressure. Pressure affects carbonate solubility in seawaterand thus Ksp and temperature. In order to depict the pressure effect alone, we prepared depth profiles of saturation
states for calcite and aragonite at St. 96 (5° 00′ N, 144°55.7′ E, water depth 4102 m, sampling date April 22,
1995) of the NOPACCS NH95-1 cruise (data wereKawahata et al. Progress in Earth and Planetary Science(2019) 6:5

a

b

c

d

Page 7 of 37

Fig. 5 Carbonate saturation state of seawater (Ωcal and Ωarag) influenced by changes in physical condition (a, c) and ongoing ocean acidification(b, d). The surface seawater compositions in the tropical western Pacific were used as an example (St. 96 of NOPACCS NH95-1 cruise; 5° 00′ N,
144° 55.7′ E; water depth 4102 m; sampling date April 22, 1995; data available from http://pacifica.pices.jp/table/). The red dotted lines in panels b
and d represent the possible future trajectory of seawater compositions, as influenced by ocean acidification (Caldeira and Wickett 2003). The chemical
equilibrium calculation was conducted using CO2calc (Robbins et al. 2010). Note that the CO2 system change in seawater caused by ocean acidification ismuch more significant to the carbonate saturation state of seawater, as compared with physical condition changes including temperature andpressure variations

Table 1 Water properties under different environmental conditions. In particular, [CO32−] and the corresponding saturation states ofcalcite and aragonite (Ωcal and Ωarag) are calculated for modern, 2110 AD, and Cretaceous conditions using CO2calc (Robbins et al. 2010).
The numbers in the gray color column show the initial settings for the calculation

Here, we assumed that the pressuere at 1 km water depth roughly corresponded to ~ 100 bar (1000 dbar)
Salinity was fined at 34

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downloaded from http://pacifica.pices.jp/table/), as shown in
Fig. 4. The pressure influence becomes dominant at greater
depth because Ksp increases 1.72 times from the surface to4000 m depth (Fig. 4d, e; Table 1). The corresponding saturation levels of the carbonate ion concentration [CO32−]sat
are 41 μmol kg−1 and 70 μmol kg−1 at the surface and4000 m depth, respectively. Note that the changes in CO2
system chemistry, i.e., degradation of OM and pH decrease,are responsible for the rapid decreases in Ωcal and Ωarag forthe depth interval below the surface layer.
Figure 5 illustrates a more quantitative comparison of
the factors influencing Ω. The surface water compositions at St. 96 in the western tropical Pacific are again
used as an example. Temperature decrease and pressure
increase result in decreases in Ωcal and Ωarag in seawater(Fig. 5a, c), but the condition of undersaturation with respect to aragonite is never reached. On the other hand,the significant changes in Ωcal and Ωarag can be causedby ongoing and future changes in the marine CO2 system, indicating that OA is the prime driver of Ω under
the conditions of climate change.

such as foraminifera, coccolithophores, mollusks, andbrachiopods with skeletons composed of low-Mg calcite(Mg/Ca molar ratio is below 4 mol%) reduce the Mg/
Ca ratio in the calcifying solution, and their skeletons
are generally more resistant to OA.
In the pelagic ocean, up to 80% of the biogenic carbonate is remineralized in the upper water column (Millero 2007). After the remaining portion settles through thewater column, the oxidation of OM at the sediment–waterinterface plays a major role in diagenetic dissolution processes (Eqs. 2 and 4 in Fig. 1). As described in the following sections, the formation and dissolution of carbonatesin the context of CO2-induced OA are examined in termsof Ω for a given environment. Although differences in theextent of carbonate dissolution in the ocean basins are primarily controlled by seawater pCO2, the use of simple
thermodynamic relationships in seawater systems is notalways an accurate predictor of biogenic carbonate dissolution due to the complex interplay of mineral surfaces
and microbial processes, the presence of reaction inhibitors, the heterogeneity of trace element distribution, differences in the susceptibility of skeletal ultrastructure, andso on (Morse et al. 2007; Cusack and Freer 2008; Iwasaki
et al. 2015; Yoshimura et al. 2017). For example, the impact of preferential dissolution is seen in a size-dependentchange of the mass fluxes of foraminifera, along withwater depth and their relevance to total carbonate flux
(Thunell et al. 1983).

Kawahata et al. Progress in Earth and Planetary Science

Characteristics of biotic and abiotic-carbonate minerals

The chemical composition, ultrastructure, and organic
components of carbonate minerals ultimately determine
their physical and chemical properties. An apparent difference in mineral solubility exists between the two major
polymorphs of calcium carbonate, i.e., calcite and aragonite (Morse et al. 2007; Ries 2011a, b). Calcite has the lower
solubility with a stoichiometric solubility product in seawater K′sp(stoich) of 6.65, whereas aragonite is more soluble(K′sp(stoich) = 4.29, Morse et al. 1980). The solubility of calcite increases with increasing Mg content. The solubilityof calcite, with an Mg/Ca of 0.14, is calculated to beequivalent to that of aragonite (Ries 2011b).
The physical properties of biominerals outperform abiotic
minerals with respect to toughness, hardness, multi-scaleordering, and architecture (e.g., Ma et al. 2009). In order tofacilitate calcification, biogenic calcium carbonate has thecompositional and spatial complexity of inorganic and organic compounds (Tamenori et al. 2014). One example isMg, an important minor constituent element in the hexagonal calcite structure (Finch and Allison 2008; Branson
et al. 2013; Yoshimura et al. 2015), which affects polymorph
mineralogy, the rate of crystal growth, and the solubility ofcalcium carbonate (Davis et al. 2000; De Groot and Duyvis
1966). The skeletons produced by anatomically simple
organisms such as corals, macrophytic algae, sponges,
and bryozoans are composed of a high-Mg calcite containing > 4 mol% MgCO3 in solid solution. In the modern marine environment, high-Mg calcite and aragonite
form via abiotic reactions. Because the presence of Mg2+
in a fluid is known to inhibit calcite crystal growth(Davis et al. 2000; Mucci and Morse 1983), certain taxa

Ocean acidification under increased pCO2 conditions in
the atmosphere

Anthropogenic CO2-driven OA has been reported to posea severe threat to marine calcifying organisms because thecalcification rates of organisms generally decrease in acidified seawater (e.g., Gattuso et al. 1999; Iguchi et al. 2012,2014; reviewed in Kleypas et al. 2006; Kroeker et al. 2013).
Kroeker et al. (2010) conducted a meta-analysis of 149
results (including 18 coral, 1 large-benthic foraminifera,
and 15 coccolithophorid studies) and reported generally
negative effects on survival, calcification, growth, andreproduction, despite significant variations among taxonomic groups.
General features of the response of coral to ocean
acidification

The work of Kleypas et al. (1999) is widely recognized as
one of the pioneering studies focusing on the decrease ofcoral calcification under the influence of increased
atmospheric pCO2, although the term “ocean acidification” was not used in this paper. Since then, experimental studies on calcification versus Ωarag at the levelof marine organisms and communities have become
more prevalent (Orr et al. 2005). In a recent review, themajority of corals were identified as showing a negative(2019) 6:5

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response to a decrease in aragonite saturation (Wittmann and Pörtner 2013).
Although some decreases in calcification rates in acidified seawater have been observed among coral species(Ries et al. 2009; Comeau et al. 2014; Iguchi et al. 2014;
Sekizawa et al. 2017), some species are insensitive to acidified seawater under certain conditions (e.g., Holcomb
et al. 2010; Rodolfo-Metalpa et al. 2010; Chauvin et al.
2011; Ohki et al. 2013; Takahashi and Kurihara 2013;
Nakamura et al. 2017). It is suggested that increased pCO2
may facilitate the onset of photosynthesis by increasing
HCO3− levels in acidified seawater (Marubini et al. 2008;
Comeau et al. 2013). Although the relationship between
coral calcification and photosynthesis by symbiotic algaein acidified seawater is not straightforward (Kleypas et al.
2006; Tanaka et al. 2007), it is evident that increased
HCO3− in acidified seawater partially facilitates coral calcification (Marubini et al. 2008; Comeau et al. 2013).
The impact of OA on corals has also been studied at
the ecosystem level. Kleypas and Langdon (2006) reported that a Ωarag of 3.3 is likely the environmental
limit for the existence of tropical coral reefs where calcification exceeds the dissolution of calcium carbonate.
Hoegh-Guldberg et al. (2007) discussed possible futurecoral reef scenarios under rapid climate change and OA
and concluded that a pCO2 above 500 ppm presented ahigh risk to coral reefs, with a potential for the loss ofcoral-dominated ecosystems. Recently, Cyronak et al.
(2018) proposed a biogeochemical approach based on
DIC–TA plots to monitor the net calcification potentialof coral reef ecosystems under anthropogenic perturbations including ocean warming and acidification.
A poleward shift of the northern limit of coral distribution along the Japanese Islands has been reported andattributed to global warming (Yamano et al. 2011). However, OA in temperate waters may influence a northward
shift of coral distribution.

cultured benthic foraminifera (Ammonia sp.) under constant DIC concentrations or constant pH conditions andreported that increased [CO32−] caused increasing
size-normalized shell weights and growth rates.
Some culturing results have suggested a new perspectiveon acidification. When we attempt to evaluate the effect of
OA on foraminifera, we have to consider (1) that the biological responses to OA differ among different life stages(i.e., gamete, neanic, or mature) and (2) the effect of eachcarbonate species (i.e., [CO32−], [HCO3−], and [CO2]) on
shell calcification. Kroeker et al. (2010) reported that thelarval and juvenile phases of marine organisms might bemore vulnerable to OA. In fact, cultured juvenile Neogloboquadrina pachyderma (sin.) individuals were more affectedby reductions in pH than adults (Manno et al. 2012).
Other culturing experimental studies using asexually
reproduced individuals have revealed different trends
among genera. Hikami et al. (2011) observed opposite responses among two reef-dwelling (large-benthic) foraminifera under different pCO2 culturing conditions, which
may have been due to the different symbiotic algae and/or
biological-calcification system. The shell weights of two
foraminifera, Amphisorus and Calcarina, were controlled
mainly by [CO32−] and pCO2, respectively (Fig. 6).
Several studies have suggested size-normalized planktic
foraminiferal shell-weight as a proxy for ambient seawater[CO32−] (e.g., Barker and Elderfield 2002; Bassinot et al.
2004; Naik et al. 2010). Moy et al. (2009) compared theshell weights of planktic foraminifera (Globigerina
bulloides) from sediment trap samples (i.e., modernday) with those in Holocene-age sediments. The
modern-day shell weights were 30–35% lower thanthose of the Holocene-age sediments in the Southern
Ocean. In addition, the fossil shell weight of this species(G. bulloides) tracks the Vostok pCO2 record over the last
50 kyr (e.g., Barker and Elderfield 2002; Moy et al. 2009).
However, at the same pressure-normalized [CO32−],
Pulleniatina obliquiloculata from the Pacific were consistently 10 μg heavier than those from the Indian Ocean(Broecker and Clark 2001). de Villiers (2003, 2004)
suggested that Ω is not the only factor controlling theshell weight of several planktic species, and Marr et al.
(2011) suggested that temperature is a prominent control on
G. bulloides shell weight (negative correlation). GonzalezMora et al. (2008) proposed both water temperature andpCO2 as the main controlling factors on shell calcification.
This was based upon the observation that G. bulloides shellweight followed the opposite trend to Mg/Ca paleothermometry and Vostok pCO2 during 250–160 kyr in the Western
Mediterranean. However, they also reported that the shellweight of N. pachyderma (dex.) was not related to either parameter, but their optimum growth condition,
Since temperature controls CO2 solubility (i.e., pCO2
and [CO32−]) in seawater, both [CO32−] and temperature

Kawahata et al. Progress in Earth and Planetary Science

Response of foraminifera to ocean acidification

Recent foraminiferal studies (both planktic and benthic)
generally exhibited negative responses to OA (e.g., Russell et al. 2004; Kuroyanagi et al. 2009; Kroeker et al.
2010; Keul et al. 2013). Culturing results showed thatshell calcification decreased under low [CO32−]. For example, culture experiments suggested that the higher is[CO32−], the heavier is the shell produced by planktic
foraminifera (Orbulina universa) (Bijma et al. 1999).
Moreover, the shell weight of Globigerinoides sacculifer
increases 2.1–3.6 μg for every 100 μmol kg−1 increase in[CO32−] (Bijma et al. 2002). Although this effect on shell
production would be species-specific, the shell weightand calcification rate of both O. universa and G. sacculifer have been shown to exhibit increases under high[CO32−] (Lombard et al. 2010). Keul et al. (2013) also

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

Page 10 of 37

a

c

b

d

Fig. 6 Comparison of an acidification condition experiment and the growth weight of foraminifera. Least mean square (± standard error) adjusted forthe rearing tank of (a) the square root of the shell weight of A. kudakajimensis and (b) the cube root of the shell weight of C. gaudichaudii after theocean acidification experiment (Hikami et al. 2011). In this experiment, foraminifera were cultured under five different pCO2 conditions, 245 (Low[pCO2], comparable to pre-industrial condition), 375 (Control-[pCO2], comparable to modern condition), 588 (High-1-[pCO2]), 763 (High-2-[pCO2]), and907 (High-3-[pCO2]) μatm, maintained with a precise pCO2-controlling technique. Arithmetic mean (± standard error) of (c) the square root of the shellweight of A. hemprichii and (d) the cube root of the shell weight of C. gaudichaudii after a constant carbonate ion ([CO3−2 = 217–218 μmol kg−1])experiment (pCO2 conditions of 354 μatm for Low-[HCO3−], 398 μatm for control, 448 μatm for High-1-[HCO3−], 486 μatm for High-2-[HCO3−], and511 μatm for High-3-[HCO3−]) (Hikami et al. 2011). Typical shapes of Amphisorus and Calcarina individuals are shown in panels (a) and (b), respectively

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

affect size-normalized foraminiferal shell weight. Therefore, further work on these two parameters should be
carried out to confirm the effect on foraminiferal
calcification.

Bleaching of corals and foraminifera in a warm and/or
high CO2 world

Major planktic foraminifera in the tropical/subtropicalocean, such as G. sacculifer, Globigerinoides ruber, andGlobigerinella siphonifera, bear photosynthetic symbiotic
algae such as dinoflagellates and chrysophycophyte(Schiebel and Hemleben 2017). Large benthic foraminifera, dwelling in warm shallow coral reefs, harbor several microalgae of dinoflagellates, diatoms, red algae,
and others as symbionts (Lee and Anderson 1991).
Many scleractinian corals that inhabit tropical/subtropical oceans belong to the class Anthozoa of the phylum
Cnidaria. They typically live in compact sessile colonies
of small individual animals (polyps) that produce hardskeleton structures of aragonite; these are important reefbuilders. They maintain a symbiotic relationship withphotosynthetic dinoflagellate. However, young polyps are
not born with dinoflagellate; rather, they obtain algae
from the surrounding water column and local sediment.
Bleaching is a potentially global threat to marine symbiotic calcifying organisms, especially corals and foraminifera. When stress-caused bleaching is not severe, bothhave been known to recover. However, if the algae loss
is prolonged and the stress continued, both will eventually disappear.

Page 11 of 37

Coral bleaching

Role of symbiosis in corals An important aspect in relation to the response of coral calcification to environmental changes is the existence of photosynthetic
symbiotic dinoflagellates, zooxanthellae (genus Symbiodinium). An association between reef-building corals
and their symbiotic algae is essential for coral survival innature (Yellowlees et al. 2008). In addition, coral calcification is facilitated by the photosynthesis of symbiotic
algae (Gattuso et al. 1999). Therefore, it is necessary toconsider the existence of symbiotic algae when we evaluate the effects of environmental changes on coral calcification. Coral calcification is closely related to coral
growth, which is an essential characteristic of their life
history; thus, a decrease in coral calcification would leadto negative impacts by causing difficulties in maintaining
coral life histories. In relation to the responses of corals
to environmental changes, the most significant
phenomenon is coral bleaching (e.g., De’ath et al. 2009)(Fig. 7).
Bleaching due to the collapse of coral–algal symbiosis
The effect of seawater temperature on corals has been
well studied because high seawater temperature is
known to be a main factor causing the collapse of coral–
algal symbiosis, which is followed by coral death(reviewed in Hoegh-Guldberg 1999). As to the cause ofcoral bleaching, it is suggested that higher temperatureinduces oxidative stress in corals. In regards to coral calcification, the increase in seawater temperature facilitates calcification within a range of mild temperatures

Fig. 7 Left panel shows coral bleaching occurring in the Sekisei Reef (24° N, 124° E), Japan, in September, 2016. Although many coral species areaffected by high temperature, it is known that massive or encrusting corals tend to survive even in mass bleaching events. The top of the rightpanel shows thermal stress sensitive species (Acropora and Pocillopora; losers) and the bottom shows stress tolerant species (Porites and Favites; winners)

(2019) 6:5

Page 12 of 37(Inoue et al. 2012). Based on the observation of skeletal
oxygen isotope (δ18O) and carbon isotope (δ13C) microprofiles in coral cores, Suzuki et al. (2003) reported thatsevere coral bleaching inhibited coral calcification of a
massive Porites colony in 1998. Also based on coral
cores, De’ath et al. (2009) suggested that increased
temperature stresses suppressed the coral calcification ofmassive Porites on the Great Barrier Reef, although it isthought that other environmental factors were likely related to the suppression. Using coral cores of Diploastrea heliopore, Cantin et al. (2010) reported that
increased seawater temperature decreased coral growth.
Laboratory rearing experiments have also clarified variations in the calcification patterns of corals in certain
temperature regimes (Anthony et al. 2008; Inoue et al.
2012). For example, Nishida et al. (2014a) reared coralnubbins of an encrusting coral, Isopora palifera, undervarious temperature treatments and found that calcification was suppressed in higher temperature treatments(> 31 °C). It is known that there are variations in sensitivity
to heat stress among species (namely, the winners and thelosers in coral bleaching; Loya et al. 2001; vanWoesik et al. 2011). Morphological (e.g., branched, flat;
Loya et al. 2001) and/or physiological variations (e.g., symbiont type, energy availability; Baird et al. 2009) among
species would contribute to different susceptibilities toheat stress. However, studies on calcification responses toseawater temperature among species remain limited.

evaluate how seawater temperature and acidification
affect coral calcification from the perspective of both thecoral host and the symbiotic algae. In the case of the calcification responses to temperature, the responses wereclearly different between aposymbiotic and symbiotic
polyps. The peak of calcification in symbiotic polyps wasaround 29 °C while that in aposymbiotic polyps wasaround 31 °C. Considering the fact that the calcificationof aposymbiotic polyps increased at 31 °C, the toleranceof the coral host to thermal stress seems to be higher, as
compared with the symbiotic condition. In the case of
calcification responses to seawater acidification, calcification tended to decrease according to the increase of
pCO2 in both aposymbiotic and symbiotic polyps; however, the calcification of adult branches was not decreased at higher pCO2 conditions (Ohki et al. 2013).
The fertilization effect of increased pCO2 on coral calcification may differ across life stages.
In addition to the advantage described above, the coral
primary polyp system has certain further advantages forevaluating the calcification responses of corals, as compared with experiments using the adult branches ofcorals. Primary polyps are very small and the calcification of primary polyps is visibly fast; therefore, a sufficient number of specimens is easily available. As such,the application of live tissue imaging to the coral primary polyp system would facilitate our understanding of
the calcification responses of corals to various environmental factors (Ohno et al. 2017a, b).

Kawahata et al. Progress in Earth and Planetary Science

Culture experiment using aposymbiotic polyps The
difficulty in understanding the complicated responses ofcoral calcification is partly attributed to the fact that coralcalcification is closely related to coral–algal symbiosis,
which forces us to consider both the coral host and thesymbiotic algae. Thus, separating the coral host and symbiotic algae under experimental conditions would be helpful
to better understand the responses of coral calcification toenvironmental changes. Acropora species release their gametes into seawater in a mass spawning event in early summer. The eggs do not include symbiotic algae but Acropora
species incorporate symbiotic algae mainly after their settlement. Thus, if we culture coral planulae and induce theirsettlement under artificial conditions using filtered seawater, which does not contain symbiotic algae, we can easily produce aposymbiotic polyps.
Furthermore, we can also induce the infection of symbiotic algae for aposymbiotic polyps by adding symbiotic
algae using an artificial method. Thus, this coral primarypolyp system enables us to experimentally separate theassociation of symbiotic algae from their coral host andalso allows us to determine which stage of the calcification process is sensitive to environmental changes(Inoue et al. 2012; Tanaka et al. 2013, 2014; Nishida
et al. 2014b). Using this system, we attempted to

Foraminiferal bleaching

Bleaching in algal symbiont-bearing large benthic
foraminifera Bleaching in algal symbiont-bearing, large
benthic foraminifera was first observed in laboratory experiments in Amphistegina spp. in the early 1980s, and it
was first noted in 1988 in field populations of Amphistegina gibbosa in the Florida Keys. Since 1991, bleaching inAmphistegina populations has been observed in subtropical reefs worldwide (Hallock et al. 2006). Bleaching inAmphistegina appears as partial “mottling” (small whitespots) that is uniformly pale or extensively white (Hallock
et al. 2006). Cytological studies indicate that bleaching inAmphistegina is caused by the loss of photosynthetic
microalgae and/or their photopigments; it is due to the digestion of deteriorating symbionts by the host (Talge andHallock 2003). Experimental studies suggest that bleaching in Amphistegina is induced by photoinhibitory stress,
in particular by the shorter (300–490 nm) wavelengths ofsolar radiation (Williams and Hallock 2004).
Temperature stress is also related to foraminiferal
bleaching. In culture experiments, Talge and Hallock(2003) found that bleaching in Amphistegina occurred
more intensively when incubated at 32 °C than at 25 °C,
indicating that bleaching was enhanced by elevated(2019) 6:5

Page 13 of 37

temperature with photic stress. Later, Schmidt et al.
(2011) demonstrated bleaching (the loss of photopigments) and decreased photosynthetic efficiency in otherdiatom symbiont-bearing foraminiferal species (Amphistegina radiata, Heterostegina depressa, and Calcarina
hispida) under high temperature stress (> 30 °C). Subsequent similar experimental studies confirmed bleaching
and reduced photosynthetic efficiency in many othersymbiont-bearing foraminiferal species under hightemperature stress (e.g., van Dam et al. 2012; Fujita et al.
2014; Schmidt et al. 2014). As summarized in Sheppard
et al. (2009), bleaching at elevated temperature is related
to the photoinhibition of algal symbionts. Temperature
limits the rate of electron transport in photosystems andthe turnover of associated D1 proteins. As a result, therate of excitation (light capture by reaction centers)
exceeds the rate of light utilization (photochemistry).
This excess energy results in a buildup of reactive oxygen species (ROS). The overproduction of ROS causes
damage to the photosynthetic apparatus of symbionts aswell as to host cells.
A meta-analysis of available experimental data up to theyear 2014 (Doo et al. 2014) showed a negative trend between growth and photosymbiont health under a + 4 °C
warming scenario, as compared with the present-day

condition (Fig. 8). Host foraminifera became inactive,
grew slowly, and tended to have increased mortality(Fig. 8a). Algal endosymbionts, irrespective of symbiont
type, exhibited (i) decreased concentrations of chlorophylls and other photopigments, (ii) a decline in the maximum quantum yields of Photosystem II (PSII; Fv/Fm)
and other photophysiological parameters, (iii) reduced
numbers of viable symbionts and high numbers of deteriorating symbionts, and (iv) decreased levels of RuBisCO
protein (the enzyme responsible for fixing CO2).

Kawahata et al. Progress in Earth and Planetary Science

Responses of symbiont-bearing benthic foraminifera
in a warm and/or high CO2 world Foraminiferal–algal
associations (as holobionts) exhibited bleaching, decreased net O2 production rates, increased respiration
rates, decreased organic carbon (C) and nitrogen (N)contents, and decreased molar C/N ratios (summarizedin Fujita et al. 2014). This negative trend in foraminiferal
responses is supported by culturing experimental studies
conducted after that of Doo et al. (2014) (e.g., Sinutok
et al. 2014; Prazeres et al. 2016; Schmidt et al. 2016).
These experimental studies suggest that combined
temperature and light stresses induce the inactivation ofphotosystems, bleaching, and reduced energy storage.
This damage to algal symbionts results in a lack of

Fig. 8 Meta-analysis of the effects of ocean warming on the growth (a) and photosymbiont health (b) of algal symbiont-bearing large benthic
foraminifera, performed by Doo et al. (2014) (the figure is modified from Doo et al. 2014). Eight publications on culturing experimental studies up
to the year 2014 were compiled for this meta-analysis, using log-transformed response ratios (LnRR) comparing the present-day condition and a
projected future scenario (Representative Concentration Pathway 8.5 from the Intergovernmental Panel on Climate Change, which projects changes of+ 4 °C by the year 2100; IPCC 2013). Negative LnRR values represent negative impacts, whereas positive values indicate positive responses. Circles
indicate the mean value, with error bars indicating 95% confidence intervals. The variability within one group comes from data between different
studies as well as those within the same study. The number of data combined per treatment (group) is shown in parentheses on the x-axis. An asterisk
indicates a significant difference compared with the present-day condition. All all foraminifera, Dino/Por dinoflagellate symbiont-bearing porcelaneous
foraminifera, Dia/Hya diatom symbiont-bearing hyaline foraminifera, Dia/Por diatom symbiont-bearing porcelaneous foraminifera, Rho/Por rhodophytebearing porcelaneous foraminifera. Foraminifera representative of each group are shown in the small images

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

movement, reduced growth, changes in symbiont genotypes, increased susceptibility to bacterial/algal infection,disease, and finally death of the host foraminifera. Thus,
bleaching and associated photophysiological parametersin algal symbiont-bearing foraminifera would be usefulas a sensitive bio-indicator of ocean warming.
Results of the meta-analysis also imply that diatom
symbionts appear more resilient than dinoflagellate symbionts under ocean warming (Fig. 8b). A meta-analysis
for OA scenarios also suggests that hyaline taxa with
diatom symbionts appear more resilient than porcelaneous taxa with dinoflagellate symbionts (Doo et al. 2014).
These contrasting differences in host growth and photosymbiont health between two foraminiferal taxa with different algal symbionts will likely result in the dominanceof hyaline foraminifera over porcelaneous foraminifera
in symbiont-bearing reef foraminiferal communities in
the future, as well as the geographical expansion of theformer taxa to higher latitudes (Weinmann et al. 2013).
The resilience of diatom symbionts over dinoflagellate
symbionts under ocean warming may also partially explain the dominance of large benthic foraminifera during
the warm period of the Cenozoic. During the Paleogene,hyaline rotaliid taxa such as Nummulites dominated anddiversified in tropical to subtropical carbonate environments (Hallock 2000). Based on the fact that their modern
descendants host diatom symbionts, these extinct hyaline
taxa possibly harbor diatom symbionts (Lee 1998). Thus, awarm, high CO2 world in the geological past was likely favored by hyaline taxa with diatom symbionts but not byporcelaneous taxa with dinoflagellate symbionts, as suggested by Hallock (2000) and Fujita et al. (2011).

Biomineralization in corals and foraminifera

In order to estimate future response of corals and foraminifera to OA, it is important to understand the basic
mechanism of small-scale calcification.

Page 14 of 37

Biomineralization of corals

Hypotheses for the mechanism of coral biomineralization
Screlactinian corals are important in terms of marine
biodiversity and the reconstruction of climate change inthe past. Since the acidification of seawater affects coralskeletons through physiological coral activities, culture
experiments using screlactinian corals under acidified
seawater have been conducted. In particular, as substantial recruitment is critical for the persistence of coral
reef ecosystems (Hughes and Tanner 2000) and primarypolyps of coral are highly susceptible to the potential effects of OA (Kurihara 2008; Suwa et al. 2010), investigations on the growth of the primary polyp skeleton havebeen conducted (Albright et al. 2008; Inoue et al. 2011;
Ohki et al. 2013). Before discussing the influence of OA,
we present the latest hypothetical models to make theprocess of biomineralization easily understandable.
Hohn and Merico (2012, 2015) classified the many
proposed ideas into four different hypotheses that reflect
alternative realizations of ion transport in corals (Fig. 9).
Hypothesis 1 (H1) states that only transcellular ion
transport of Ca2+ and HCO3− into the calcifying fluid induces coral calcification (Allemand et al. 2004) (Fig. 9a).
Hypothesis 2 (H2) proposes that Ca2+ and carbon aresupplied only by the paracellular pathway (Gagnon et al.
2012). In hypothesis 3 (H3), protons are removed fromthe calcifying fluid to increase aragonite saturation by aproton pump while Ca2+ is transported into the calcifying fluid passively via the paracellular pathway, known
as the proton flux hypothesis (Jokiel 2011; Ries 2011a).
In hypothesis 4 (H4), a combination of transcellular andparacellular ion transport can induce coral calcification(Cohen and McConnaughey 2003; Nakamura et al.
2017). Hohn and Merico (2015) examined the plausibility of different calcification mechanisms in relation to
OA using a mathematical model. The results suggestthat H4 and a combination of the processes in H1 andH2 are consistent with changes in calcifying fluid pH

Fig. 9 Different ion transport hypotheses tested in this study and their simulation results (Hohn and Merico 2015). (H1) transcellular ion transport
only, (H2) paracellular ion transport only, (H3) proton flux hypothesis, and (H4) combination of transcellular and paracellular ion transport. The
upper panels visualize the conceptual models of the different transport hypotheses (H1–H4). Ca2+ is transported into the cell through the Ca2+-channel, whereas it is transported into the calcifying fluid from the cell by the Ca-ATPase pump(2019) 6:5

Page 15 of 37

and Ca concentration and with calcein staining experiments. An efflux of Ca2+ and CO32− from the calcifying
fluid is generated by the paracellular pathway, which results in a leakage of ions that counteracts the concentration gradients maintained by the transcellular pathway.
Future ocean acidity would exacerbate this carbonate
leakage and reduce the ability of corals to accrete
CaCO3. However, this result was obtained by computer
simulation and was not confirmed by observations infield and culture experiments. Ca isotopic work could
provide some clarification on this issue.

calcification has been estimated (Hohn and Merico2015), indicating the importance of transcellular transport, which is always directed into the calcifying fluid.
However, there also exists carbonate leakage, which appears to yield greater difficulty in accreting calcium
carbonate.
Inoue et al. (2015) further investigated the relationshipbetween the isotope fractionation of Ca in coral skeletons and the environment (temperature, seawater pH,
and light intensity) using cultured corals. As a result,only temperature affected δ44Ca in Porites corals and its
temperature dependence was 0.02‰/°C. Interestingly,
negligible changes in isotopic fractionations of Ca were
found in the pH and light intensity experiments,
whereas significant differences in skeletal growth rate
were observed in both experiments. Therefore, thetransport path of Ca2+ and its activity might have aminor impact on skeletal growth; instead, carbonate
chemistry in the fluid would predominantly control the
precipitation rate, as suggested by Hohn and Merico(2015) (Fig. 9).

Kawahata et al. Progress in Earth and Planetary Science

Biomineralization of corals based on analyses of Ca
isotopes It has been suggested that acidified seawatersignificantly decreases the growth rate of primary polyps,
as reported for adult corals. However, recent studies suggest that corals regulate their internal pH and thencounteract any acidification (Ohno et al. 2017a; McCulloch et al. 2012). In fact, the mechanism by which corals
control the transport of Ca2+ and other ions from seawater are largely unknown. Corals build their skeletonsusing extracellular calcifying fluid located at the tissue–
skeleton interface.
In terms of coral calcification, Ca2+ and CO32− must be
constantly supplied to the calcifying fluid to precipitate
aragonite skeletons. It has been suggested that Ca-ATPase
controls the saturation state of aragonite in the calcifying
fluid as the pump transfers Ca2+ in exchange for H+ acrossthe calicoblastic cells leading to high calcification rates in
the calcifying fluid (Fig. 9; Al-Horani et al. 2003; Cohen
and McConnaughey 2003). Although there is no transporter known for carbonate, carbon can cross the coralcells via free diffusion of CO2 over cell membranes. Then,the CO32− ion can be resupplied because once CO2 or
HCO3− reaches the calcifying fluid, a new chemical equilibrium between CO2, HCO3−, and CO32− is established.
These pathways are called transcellular pathways since
Ca2+ and carbon must pass through the coral cells.
However, there are several reports suggesting that seawater may also enter the calcifying fluid directly, mainly
based on experiments using a membrane impermeable
dye (Cuif et al. 2011; Gagnon et al. 2012).
In order to estimate the transport of Ca2+ by the
Ca-channel and Ca-ATPase, Ca isotopes (44Ca/40Ca) inscrelactinian corals (Acropora sp. and Porites australiensis) have been measured and + 0.4‰ offset compared toinorganic aragonite was found (Böhm et al. 2006; Inoueet al. 2015) (Fig. 10). The offset suggests that Ca2+ istransported into calcifying fluid mainly via a transcellular pathway (Fig. 10), although this does not imply that
there is no direct transport (paracellular pathway) ofother ions from seawater to the calcification site. As
mentioned previously, the relative contribution of ion
transport (transcellular or paracellular pathways) to coral

Biomineralization of foraminifera

Porcelaneous imperforate and hyaline perforate testsof foraminifera Foraminiferal regulation of Ca2+ and
CO32− uptake into calcareous tests are of great interest toscientists studying OA and its history. The calcareous foraminifera consist of two groups, porcelaneous (miliolid)
and hyaline, according to their test structure. Porcelaneous species produce calcite in the form of needles with alength of a few μm in cellular vesicles (Parker 2017). The
walls of these species exhibit an opaque appearance likeporcelain. Hyaline species precipitate chamber walls extracellularly on primary organic sheets. These are constructed prior to calcification as a new chamber thatprovides nucleation sites for the initial calcite precipitation(as reviewed by Erez 2003). The crystallography of theporcelaneous and hyaline species differs from each other,as do their calcification pathways. The porcelaneous species have high Mg-calcite (Toyofuku et al. 2000; Maeda
et al. 2017), whereas most hyaline species produce low
Mg-calcite. The variability in Mg level can be explained bydifferences in the cytological pathway of Ca uptake (deNooijer et al. 2009a). Porceallaneous species form imperforate shells, whereas hyaline species form perforate shells(Ottway et al. 2018).
Since their solubility differs depending on crystallography, the responses of porcelaneous and hyaline species
to OA differ markedly. As a result of culture experiments with large foraminifera, the tests porcelaneous
species were dwarfed, whereas hyaline species showedno uniform response among the species (Kuroyanagi
et al. 2009; Fujita et al. 2011; Hikami et al. 2011). Among
recent foraminifera, Hoeglundina and Ceratobulimina

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

Page 16 of 37

Fig. 10 Schematic diagram of coral calcification in terms of Ca2+ transport. δ44Ca values of 1.88‰ and 0.73‰ are reported for seawater (Hippler
et al. 2003) and coral skeletons (as the mean value of all data) (Inoue et al. 2015), respectively. The value of − 1.2‰ represents the typical fractionation
at the Ca-channel (Wiegand et al. 2005; Böhm et al. 2006; Gussone et al. 2006). Typical concentrations of Ca2+ in seawater, calicoblastic cells, andcalcifying fluid are presented based on the reports of Al-Horani et al. (2003) for seawater and the calcifying fluid, and of Carafoli (2004) and Saris andCarafoli (2005) for the cells

are known to have aragonitic tests (Bandy 1954). Sincearagonite has greater solubility than calcite, it is presumed to be strongly affected by acidification; however,there are apparently no studies to confirm this hypothesis. Evaluation of the influence of OA on these aragonitic species is a future research need.
Visualization of extracellular pH distribution tounderstand foraminiferal calcification Based on isotope labeling experiments, calcification by foraminifera ishypothesized to involve intracellular storage of Ca2+ ionsand inorganic carbon (ter Kuile et al. 1989; Erez 2003).
Foraminifera are also considered to manipulate pH (deNooijer et al. 2009a; Bentov et al. 2009; Toyofuku et al.
2017), uptake Ca (Erez 2003; Toyofuku et al. 2008), and/or control Mg (Bentov and Erez 2006) via cytological
functions. The elemental and isotopic composition of foraminiferal calcareous tests has also constrained models

for calcification (Elderfield et al. 1996; Bentov and Erez2005). Recently, visualization methods using variablefluorescent indicators have become powerful tools for revealing the biomineralization process of foraminifera(Toyofuku et al. 2008; de Nooijer et al. 2008, 2009a, b;
Bentov et al. 2009; Ohno et al. 2016; Toyofuku et al.
2017). The variety of results and partly contradicting hypotheses and models for foraminiferal calcification makeit essentially impossible to predict their response to OA
and enrichment of dissolved CO2 (de Nooijer et al. 2014).
Indeed, the foraminiferal responses to OA are variedamong species (Kuroyanagi et al. 2009; Fujita et al. 2011;
Hikami et al. 2011; Keul et al. 2013).
Fluorescence observation has revealed that management of protons is important for foraminiferal calcification and for corals (e.g., Cai et al. 2016). The hyaline
benthic species, Ammonia sp., show a large pH variation(~ 2 units) between the inside and outside of the site of

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

calcification during the chamber formation process(de Nooijer et al. 2009a; Glas et al. 2012; Toyofuku et al.
2017). Visualization of the extracellular pH distribution
around individuals shows that chamber formation is accompanied by a strong decrease in external pH from ambient to ~ 7 (Fig. 11). Meanwhile, the highest pH reaches
more than 9 within the site of calcification. This decreasein external pH was observed over a wide range of pCO2
(from 460 up to 9010 μatm /from pH 8.0 to 6.8) and the
reduction in pH compared to that of the ambient seawater
was relatively constant over the experimental conditions.
The observed decrease in pH outside the site implies thatthese foraminifera actively pumped out protons via aphysiological process, with the flux independent of initialexternal pCO2/pH.
Under observation in the presence of a proton pumpinhibitor (Bafilomycin A1), the external pH gradient disappeared (Toyofuku et al. 2017). Further, the specimens
could no longer accomplish calcification. It is strongly
suggested that proton transport is essential to the foraminiferal calcification process. Perhaps this is due to the
conjugation of protons and Ca2+ exchange by an antiporter. These findings also suggest that the site of calcification of this species is well isolated from the ambientenvironment. Recently, these considerations have been
supported by SEM observation on the site of calcification of Ammonia (Nagai et al. 2018; Fig. 12). The coverage of a thin layer (~ 1 μm) of a pseudopodial sheet
extending over the entire specimen may be a strategy for
maintaining oversaturation close to the chamber surfaces. As a consequence, the site of calcification may extend toward the pre-existing chambers and produce athin layer of calcite over the pre-existing chambers to

Fig. 11 Planar variability in pH around the calcifying specimen ofAmmonia sp. in seawater with the pH indicator, pyranine. The gradient inpH, which decreases toward the specimen, is mainly caused by proton
diffusion away from the precipitation site. An image of foraminifera issuperimposed on the pseudocolor pH image. Scale bars indicate 100 μm

Page 17 of 37

Fig. 12 Proton pumping-based model of rotaliid foraminiferal
calcification at the enclosed site of calcification (SOC). a Magnified
SEM image of the cross section of SOC. Upper side is the outer side,
whereas the lower side is the cellular side. b Black rectangle indicatesthe observed newly calcifying chamber in (a). c Considerable
calcification model based on Toyofuku et al. (2017) and Nagai et al.
(2018). During chamber formation on an organic template (termed
the primary organic sheet, POS), the organic layers isolate the growing
calcite surface from the surrounding seawater. The chemicalcomposition of the calcification fluid of the SOC is characterized by
active, outward proton pumping coupled with an inward calcium
ion-flux. The transported proton reduced pH around the foraminiferal
microenvironment. Inorganic carbon speciation is shifted by this lower
pH, thereby increasing pCO2 of the surrounding seawater near the siteof calcification. The large gradient in pCO2 around the individual resultsin the diffusion of CO2 into the site of calcification. This inbounded CO2
reacts to form carbonate ions due to the high pH at the site of
calcification. This transported carbon source sustains CaCO3
precipitation by reacting with the inward- transported calcium ionsconstruct a unique multi-lamellar structure (Reiss 1958;
Bé and Hemleben 1970; Grønlund and Hansen 1976).
Proton pumping in the calcification of Ammonia sp.
Ammonia sp. are widely distributed in many coastal

(2019) 6:5

Page 18 of 37

environments, including tidal flats, ports, and brackishlakes. Because of its resistance to a wide range of environmental variables, it is a genus popularly employed inlaboratory cultures. It is also used to study geochemicalsignatures as a function of physical and chemical conditions. Further, Ammonia sp. serve as model species,
whose chamber wall structure resembles that of many
Rotalid species. The calcification of Ammonia sp. ischaracterized by strong gradients in pH and/or pCO2
between the site of calcification and the foraminiferal
microenvironment. The results suggest that foraminiferal calcification would be regulated by DIC concentration rather than the calcite saturation state (Fig. 12).
This uncoupling of saturation state and calcification explains the moderate response of many foraminifera to
OA and the occurrence of calcareous foraminiferal communities at undersaturated conditions with respect tocalcite but at high DIC concentrations (Keul et al. 2013;
Flako-Zaritsky et al. 2011). The reduction in oceanic pH
by increased oceanic CO2 uptake is relatively small compared with the pH decrease in the microenvironment
during Ammonia’s calcification. Hence, a relatively moderate decrease in pH may not impair foraminiferal calcification. Proton pumping is found to be the fundamental
strategy for a variety of marine calcifiers (Ries 2011a).
The high internal pH and large internal–external pH
gap associated with foraminiferal calcification predicts
that they may well produce more carbonate ions at thesite of calcification even under elevated pCO2.

In the deep sea, the degree of seawater acidity also increases with the age of the water mass because of accumulating CO2, nitrate, and phosphoric acid derived fromdecomposed OM. Therefore, the depth of the transitionzone is essentially influenced by deep water properties via
thermohaline circulation and it becomes shallower as theage of the deep water increases (Broecker and Takahashi
1977). Actually, the sedimentary carbonate content in the
Atlantic Ocean is generally higher than that of the Pacific
Ocean. In particular, the carbonate content is almost zeroin the North Pacific below 4000 m. Most carbonate isdissolved in the deep sea in the North Pacific regardless ofthe high carbonate export flux (e.g., Kawahata 2002). Thechange in depth of the transition zone must be highly correlated with DIC and alkalinity contents in the global
ocean (Figs. 2 and 4). Therefore, the dissolution and preservation of calcite play an important role in the global
carbon cycle.

Kawahata et al. Progress in Earth and Planetary Science

Deep-sea carbonate dissolution and progressive foraminiferal
shell dissolution

A more severe influence of OA can be found in the deepsea, where carbonate tends to dissolve due to its higher
solubility. This is controlled by the solubility product,
which increases with depth (pressure), and by the increased acidity caused by the degradation of OM (Figs. 1
and 4, Table 1).
Spatial and vertical heterogeneity in carbonaceous sedimentsin the deep sea

The distribution of carbonate in marine sediment is inhomogeneous (e.g., Archer 1996). For example, regionswith high carbonate content are mostly in the shallow
oceanic ridges, whereas the deep ocean basins are almost
completely devoid of carbonate due to its dissolution as aresult of the high pressure, high [CO2], low temperature,and hence, low saturation state (Eq. 4 in Fig. 1). These tworealms are separated by a transition zone spanning several
hundreds of meters in water depth, over which the carbonate content drops toward zero. The upper limit of thetransition zone is termed the lysocline and lower limit of
the transition zone is termed the carbonate compensation
depth (CCD) (Fig. 4).

Conventional proxies for carbonate dissolution

Several proxies for carbonate dissolution have been proposed based on planktic foraminiferal shells. Berger(1968) proposed a dissolution index based on empiricallyranked species of planktic foraminifera because they havea variety of morphologies and some of them are fragile
and prone to dissolution. Similarly, benthic foraminifera
normally have thick and smooth shells that are more resistant to dissolution than those of planktic foraminifera.
Therefore, the planktic to benthic foraminiferal ratio in
the sediment has been suggested as a dissolution proxy(Peterson and Prell 1985). However, the results depend on
the initial makeup of the foraminiferal population in thesediment at the regional scale. The ratio of foraminiferal
fragments to whole shells has been widely used as a calcite
dissolution index (Peterson and Prell 1985). The advantageof the fragment ratio is its simplicity; however, its determination is often subjective. These methods are not calibrated against present-day pressure-normalized [CO32−].
The size-normalized shell weight of planktic foraminifera, which was first proposed by Lohmann (1995), is themost widely used method to reconstruct past [CO32−].
Shell wall thickness and shell bulk density can be estimated using this method. A linear relationship between
size-normalized shell weight and bottom-water calcite saturation state has been suggested (Broecker and Clark
2001, 2002). This method is based on an empirical calibration of bottom water [CO32−] or Ωcal, with the average
weight of clean shells of selected species picked from narrow size ranges in core-top samples (e.g., de Villiers 2005).
However, Barker and Elderfield (2002) showed that thickness of the foraminiferal shell walls varies with growth
conditions. Spero et al. (1997) found strong correlations
between shell weight and both water temperature and[CO32−]; the warmer the water and the higher its [CO32−],(2019) 6:5

Page 19 of 37

the thicker the shells. Therefore, size-normalized shellweight techniques have uncertainties owing to the initialshell weight variation depending on Ω in the surface
water.

Using changes in the CT number histogram with theprogression of shell dissolution, a quantitative proxy thatallows quantitative measurements of carbonate loss fromforaminiferal tests prior to carbonate dissolution wasprovided, which is not influenced by initial shell condition. Selective dissolution of the inner calcite layer andchanges in the CT number histogram were also observed in G. bulloides shells from marine sediment samples (Fig. 13c, d), suggesting that the dissolution proxybased on XMCT scanning might be applicable to estimate
the amount of carbonate dissolution at the deep-sea floor.
Furthermore, characteristics of the shell dissolution
process, accompanied by changes in the micro-scale crystalline structure, should be applicable to various species
other than G. bullloides. Although several problems remain to be addressed (e.g., the time required for scanningand the imperfect understanding of the dissolution mechanism in the water column and in deep-sea sediments), a
dissolution proxy utilizing the CT number histogram
based on XMCT scanning has the potential to be the first
viable method for accurately estimating the actual amountof calcite in seawater [CO32−].

Kawahata et al. Progress in Earth and Planetary Science

Developing proxies for carbonate dissolution using X-ray
micro-computed tomography analysis

Recent progress in X-ray micro-computed tomography(XMCT) has allowed researchers to observe the internal
structures of foraminiferal shells and to evaluate shell
dissolution intensity. Johnstone et al. (2010) used an
XMCT scanner to observe the dissolution of foraminiferal shells from core top sediments and suggested thatforaminiferal shells dissolve from the inner chamber
walls and that the initial dissolution cannot be observed
from outside of the shells. Furthermore, they divided thedissolution process of foraminiferal shells into five stages
and proposed a new empirical dissolution index, XDX,
based on the appearance of the CT images, and they
provided a calibration to estimate ΔM (the loss of test
mass due to dissolution) from XDX.
In addition, XMCT makes it possible to quantitatively
estimate shell density, and to observe a high-resolution
density distribution in an individual shell. Iwasaki et al.
(2015) used an XMCT scanner to estimate foraminiferal
shell density and performed time series dissolution experiments with tests of planktic foraminifera, G. bulloides, inpH-controlled undersaturated seawater in order to observe the initial stages of the dissolution process of G. bulloides tests (Fig. 13). They employed an XMCT system(ScanXmate-D160TSS105/11000, Comscantecno Co. Ltd.,
Kanagawa, Japan) to investigate the internal structure anddensity distribution of the shells. A high-resolution setting(X-ray focus spot diameter of 0.8 μm, X-ray tube voltageof 80 kV, detector array size of 2000 × 1336, 1500
projections/360°, 2.5 s/projection) was used for threedimensional (3-D) quantitative densitometry of small foraminiferal shells. The CT number, which is the normalized
value of the calculated X-ray attenuation coefficient of avoxel in a computed tomogram, was used as an indicatorof calcite density. In their study, comparison of CT and
SEM images of the cross section of the outermost chamber of G. bulloides revealed that the inner calcite layer has
a porous microgranular crystalline structure (low-density)and the outer calcite layer has a euhedral crystalline structure (high-density).
From their time series dissolution experiment, dissolution of the G. bulloides shell started from the central
area of the shell and then spread to the porous inner calcite layer. In contrast, the outer calcite layer with a euhedral crystalline structure was resistant to dissolution
and was well preserved (Fig. 13a, b). Such selective dissolution of the porous inner calcite layer was also evident as changes in the histogram of the CT numbers.

Response of marine calcifiers to different environments

The Earth’s surface has previously experienced a highpCO2 environment, which was generally associated with“hot house” conditions. Carbonate in the deep-sea wasabruptly dissolved at the Paleocene/Eocene boundary
due to severe OA. However, especially in the warm Cretaceous, abundant carbonate precipitated at the seafloor.
These observations clearly demonstrate that high pCO2
alone did not always result in OA.
Extreme warmth in the Cretaceous greenhouse system

The mid-Cretaceous is widely accepted as the archetypalice-free greenhouse interval in Earth’s system history. To
understand the greenhouse climate system, the Cretaceous paleoclimate has been widely discussed in relation
to geological, geochemical, and paleontological records.
Since early reports of mid-Cretaceous tropical andsubtropical plants from the northern high latitude(e.g., Nathorst 1911), many paleontological studies
have described a warm climate even at high latitudes(e.g., Tarduno et al. 1998). This warm climate is attributable
to high pCO2 and possibly methane levels, in association
with the high production rate of oceanic crusts, enhanced
continental rifting, formation of large igneous provinces(LIPs), and metamorphic decarbonation (although volcanic
activities releasing volcanic dust sometimes act as cooling
events) (e.g., Larson 1991; Kuroda et al. 2007; Moriya 2011;
Scudder et al. 2016; Brune et al. 2017).
The extensive warmth observed during the
mid-Cretaceous is also supported by the oxygen isotopic
paleothermometry of well-preserved planktic foraminifera

Kawahata et al. Progress in Earth and Planetary Science

a

b

c

d

Fig. 13 (See legend on next page.)(2019) 6:5

Page 20 of 37

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

Page 21 of 37(See figure on previous page.)
Fig. 13 X-ray micro-CT cross-section images and histogram of CT number for G. bulloides shells: a original fresh shell before dissolution and b a
partly dissolved shell after a low-pH condition in a 6-day dissolution experiment. c Relatively preserved shell and d relatively dissolved shell collectedfrom the surface sediment at a water depth of 1362 m and 3135 m of the North western Pacific, respectively. As dissolution progressed, the inner sideof the foraminiferal shell was selectively dissolved, and the mode of the CT number histogram gradually decreased. This CT image was obtained byscanning using a X-ray micro-CT system (ScanXmate-D160TSS105/11000, Comscantecno Co. Ltd., Kanagawa, Japan) at the Tohoku University Museum.
Data are cited from Iwasaki et al. (2015)(so-called glassy foraminifera) (e.g., Wilson et al. 2002;
Moriya et al. 2007; Bornemann et al. 2008; Friedrich et al.
2008). However, the Cretaceous equatorial sea surface
temperature (SST) estimated in twentieth-century studies
was less than 25 °C, which is much lower than it should
be and even lower than the modern equatorial
temperature. The highest SST of 29 °C in the modern
ocean is observed in the Western Pacific Warm Pool
(WPWP) (Fig. 14). From the surface, water temperature
decreases vertically downward to ~ 1 °C at the bottom of
the modern equatorial Pacific, with a steep thermocline at
100–300 m, indicating that the water column is wellstratified in the modern equatorial ocean. Compared with
the SST in the WPWP in the last glacial maximum, avalue of ~ 25 °C in the Cretaceous greenhouse world issurprisingly cool. This cool equatorial condition in thegreenhouse world is termed the Cool Tropic Paradox(D'Hondt and Arthur 1996).
At the beginning of the twenty-first century, Pearson
et al. (2001) found that early diagenesis on or just under
the sea floor lowers the isotopic temperatures recorded inforaminiferal fossils. While many authors in the twentiethcentury utilized foraminiferal fossils with minor
recrystallization for isotopic thermometry, new data generated from extremely well-preserved “glassy” foraminifera
conclusively indicated that SSTs of the Cretaceous greenhouse were extremely high. Isotopic records show that
mid-latitude SSTs of ~ 25 °C were higher than the modern
mid-latitude temperature (Fig. 14). The difference between
the modern and Cretaceous SSTs becomes prominent in

45

Foraminiferal calcite (Hedbergella)

40

Foraminiferal calcite (globular planktics)
Vertebrate phosphate

35

Highest temperature in modern WPWP

30
25

Mid-latitude temperature in modern Pacific 37˚N

15
30

15

45

10
Low latitude temperature in modern Pacific 60˚N

5

60
75

Paleolatitue (˚)

0
20

90

0
Maa.

70

Campanian Sa. Co. Tur. Cenom.

80

90

100

Albian110

Aptian120

Barr. Ha. Valang. Berri.

130

140

Age (Ma)
Fig. 14 Evolution of the Cretaceous ocean temperature estimated from oxygen isotopic records. Circles, triangles, and squares represent datafrom Hedbergella (planktic foraminifera), globular form planktic foraminifera, and surface dwelling vertebrate remains, respectively. Colors for eachsymbol indicate the paleolatitude where the data were obtained. While foraminifers were recovered from deep-sea cores, vertebrate remains occurredin outcrops, implying that vertebrate remains may not record open ocean temperatures. δ18OSW is estimated as − 1 (‰, VSMOW) (Shackleton and
Kennett 1975). δ18OSW is adjusted for latitudinal gradient in surface water δ18OSW (Zachos et al. 1994). Isotopic temperatures were calculated using theequations of Bemis et al. (1998) for foraminiferal calcites and Pucéat et al. (2010) for vertebrate remains. Paleolatitudes were computed by the
Paleolatitude Calculator for Paleoclimate Studies (version 2.1) (van Hinsbergen et al. 2015) if they were not provided in the literature. Data are citedfrom Huber et al. (1999, 2002), Bice et al. (2003, 2006), Moriya et al. (2003; 2007; 2011), Pucéat et al. (2003, 2007), Forster et al. (2007), and Bornemannet al. (2008). Maa. Maastrichtian, Sa. Santonian, Co. Coniacian, Tur. Turonian, Cenom. Cenomanian, Barr. Barremian, Ha. Hauterivian, Valang. Valanginian,
Berri. Berriacian, WPWP Western Pacific Warm Pool(2019) 6:5

Page 22 of 37

the Cenomanian–Turonian (~ 95 Ma), which appears tohave been the warmest interval on Earth for at least the last
100 million years (e.g., Wilson et al. 2002). Backgroundequatorial SSTs in this interval were ~ 35 °C and the highest SST exceeded 37 °C, indicating the cirrus cloudnegative-feedback hypothesis proposed for the modern
ocean might not be valid, or that it had another steady statein the ice-free greenhouse system. Although SSTs became
cooler in the late Cretaceous, equatorial and mid-latitude
SSTs were still higher than those of the modern ocean.
The other notable disparity was a significantly weak
meridional SST gradient in the Cretaceous. This was especially obvious in ~ 90 Ma (Fig. 14). While the Cretaceous SST gradient between the equator and 60° in
latitude is ~ 14 °C, the gradient in the modern ocean is~ 25 °C. This observation indicates enhanced heat transport from the equator to the poles in the Cretaceous
greenhouse system. However, since numerical examinations (e.g., general circulation models) cannot simulate
this weak SST gradient with reasonable pCO2, unknown
heat transporting mechanisms in the greenhouse systemor high concentrations of other greenhouse gases, suchas methane, are expected (Bice et al. 2006). On the other
hand, fully coupled models (e.g., CCMS [Community
Climate System Model], MIROC [Model for Interdisciplinary Research on Climate], etc.) have been addressing
this important problem. Upchurch et al. (2015) first succeeded in describing the weak meridional SST gradient
without unreasonable artificial tuning. This significant
advancement is opening the door to understanding thegreenhouse climate system in collaboration withmulti-site and -depth proxy archives.

survival of planktic foraminifera under low-oxygen concentrations remain obscure.
Culture experiments have examined the direct effects
of dissolved oxygen concentration on planktic foraminifera. Kuroyanagi et al. (2013) cultured planktic foraminifera (both symbiont and non-symbiont species) under
different oxygen levels. These foraminifera can survive,
add chambers, feed, and undergo gametogenesis evenunder dysoxic conditions (~ 0.7 mg O2 L−1 or ~ 22 mmol
O2 m−3). Furthermore, the gametogenesis rate and time
to gametogenesis did not show any obvious trends
across oxygen levels (Fig. 15). On the other hand, otherresults from the culture experiments showed that noplanktic foraminifera could survive for more than 48 hin the presence of H2S (~ 2 and ~ 9 mg H2S L−1)
(Kuroyanagi et al. 2019). Thus, critical boundary conditions for the survival of planktic foraminifera under
low-oxygen concentrations are still unknown due to thedifficulty of maintaining extremely low oxygen levels
throughout a culture. Several studies have reportedchanges in oxygen-minimum zone conditions corresponding to planktic foraminiferal disappearance in Quaternary
sediments (Fenton et al. 2000; Reichart et al. 2004). Undera “business-as-usual” scenario, modeling results have revealed a dramatic 50% increase in the suboxic water volume due to more stratified surface water with higher CO2
levels until 2100 AD (Oschlies et al. 2008). Therefore, culturing planktic foraminifera under suboxic water conditions (0 to ~ 20 mmol O2 m−3) might provide essential
information and contribute to estimating changes in the
future global carbon cycle.

Kawahata et al. Progress in Earth and Planetary Science

Stratified ocean in the Cretaceous greenhouse system

Warmth in the middle Eocene climatic optimum event andforaminiferal photosymbiosis

It is well known that in the mid-Cretaceous, the greenhouse period, oceanic anoxic events (OAEs) occurred
several times (e.g., Jenkyns 2010) due to the lower solubility of oxygen into the seawater at higher SST andlower ventilation of the water mass. One of the most interesting features of the anoxic sediments is that large
quantities of organic carbon were often deposited in association with carbonates (e.g., nannoconid, foraminifera). In fact, well-preserved planktic foraminifera havebeen discovered from the black shale collected fromODP Site 1049 in Tethys (Norris and Wilson 1998), indicating that carbonate was not significantly dissolved
even in suboxic or anoxic waters.
The oceanic redox state is a critical determinant in theevolutionary history of life on Earth (Raup and Sepkoski1982). High turnover rates of fossil planktic species(planktic foraminifera and calcareous nanofossils) andradiolaria dwelling in the mid-water column have been
observed at or near these OAEs (Leckie et al. 2002; Itaki
2016), although the critical boundary conditions for the

The Eocene epoch is characterized by overall warmth,with a gradual cooling trend and short-term hyperthermal events. The SST at low latitudes was estimated to beas high as 30–35 °C (Pearson et al. 2007). Especially during the extreme warmth of the middle Eocene climatic
optimum (MECO) event, pCO2 increased by a factor of
2 to 3, and SST rose by 3–6 °C to exceed 28 °C even insouthern high latitudes (Bijl et al. 2010). Such conditions
might be disastrous to photosymbiotic organisms basedon our current knowledge on modern coral “bleaching,”
i.e., a collapse of photosymbiosis, via heat stress. It is
pointed out that symbiont “bleaching” might have also
occurred in planktic foraminifera.
The photosymbiotic ecology of planktic foraminifera
can be detected even from fossils with stable isotopic
signatures of symbiont photosynthesis recorded in foraminiferal tests (Oppo and Fairbanks 1989; Norris 1996).
A characteristic carbon isotopic signature has been used
as an indicator; the δ13C value of symbiont-bearing species is more 13C-enriched than the non-symbiotic

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

Page 23 of 37

Fig. 15 Gametogenesis rates (left) and days to gametogenesis (right) in Orbulina universa (solid red line) and Globigerina bulloides (broken blueline) at different dissolved oxygen (DO) saturation levels (figure modified from Kuroyanagi et al. 2013). Although DO environments are classified
differently in different studies (e.g., Jorissen et al. 2007), a DO of 10% (~ 0.7 mg O2 L−1 or ~ 22 mmol O2 m−3) was generally classified as “dysoxic”

species, reflecting a selective incorporation of 12C during
symbiont photosynthesis (“photosymbiotic effect”); thegreater the symbiont photosynthesis, the larger the effect. Therefore, the effect is enhanced as the host grows
its body size and progressively increases the number ofthe symbionts. This results in a positive relationship between test size and δ13C. Based on this concept, many
authors have studied the size-related δ13C profiles of fossil foraminiferal tests of extinct species to identify theirphotosymbiotic ecology (e.g., Norris 1996; Bornemannand Norris 2007; Birch et al. 2012). Moreover, the δ13C/
size positive gradient has been used as an indicator of
the intensity of the photosynthetic activity of symbionts(Kelly et al. 1996; Wade et al. 2008; Edgar et al. 2013).
The δ13C/size positive gradient of some photosymbiotic species deteriorated during the peak MECO conditions, as compared with pre-event and post-event levels(Edgar et al. 2013). This deterioration could be attributed to symbiont “bleaching” during the higher SSTs in
the peak MECO period. Despite these possible crises ofphotosymbiosis, many Eocene planktic foraminifera arethought to have been photosymbiotic (Pearson et al.
2006). The chamber-by-chamber isotopic profiles ofthree representative Eocene genera, Morozovella, Acarinina (both known as symbiotic), and Subbotina (known
as non-symbiotic), recovered from IODP Site U1407 exemplify these ecologies. As expected, the results revealeda clear contrast in δ13C profile between the symbiotic
and non-symbiotic species (Fig. 16a, b).

A notable finding here is that the last chamber of thetwo symbiotic species showed remarkably low δ13C values,
comparable with that of the non-symbiotic Subbotina.
This may indicate that at the time of calcification of thelast chamber, these two symbiotic species became virtually“non-symbiotic” with the loss of their symbionts. Thismay have been caused by the digestion or dissolution of
the symbionts, as seen in cultured modern photosymbiotic
species (Bé et al. 1983; Takagi et al. 2016a). The modern
and fossil examples of symbiont loss mentioned above
should be part of the natural life processes of foraminifera.
Such lack of a “photosymbiotic effect” in only the lastchamber cannot be identified using conventional δ13C/size
methods. However, it can be distinguished inchamber-by-chamber δ13C profiles (Fig. 16a). Therefore, achange in photosymbiotic activity, or even a “bleaching”
phenomenon, can theoretically be identified by precise
δ13C analyses of individual chambers.
Although the possibility of “bleaching” in the Eocene
warm period cannot be denied, caution should be used
when interpreting the δ13C profiles. Following the rule ofmetabolic rates, an elevation of SST enhances both photosynthetic rates and respiration rates with different
temperature sensitivities. Temperature sensitivity is higher
for respiration than for photosynthesis (e.g., Lombard
et al. 2009). This means that deterioration of the δ13C gradient can also be caused by elevation of the respiration
rate under high SST conditions because the effect of respiration on test δ13C shows the opposite trend. When

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

aPage 24 of 37bFig. 16 Comparison of isotopic results for Eocene species and modern cultured photosymbiotic species. a Chamber-by-chamber δ13C versus
cumulative test mass of Eocene species. The same shade is used for symbols of the same species (black: Morozovella crater; gray: Acarinina
cuneicamerata; white: Subbotina eocaena). Data for the same individual are connected in ontogenetic order. Error bars represent analytical errors(± 0.1‰). Each individual test was dissected into chambers using a micro-blade (e.g., Takagi et al. 2015; 2016b), then analyzed by a customized
continuous-flow mass spectrometry system that can measure micro-volume carbonate as small as a single chamber (Ishimura et al. 2004, 2008).
Symbiotic species Morozovella and Acarinina tended to show slight positive shifts until the penultimate chamber and then the final chamber showed
drastically more 13C-depleted δ13C (values 0.5–1.0‰). In contrast to these species, Subbotina showed comparatively 13C-depleted δ13C values throughits ontogeny, which is in good agreement with the known non-symbiotic nature of this species. b Growth and test isotopic results for the modern
symbiotic species Globigerinoides sacculifer cultured in a laboratory (Takagi et al., 2018). Chlorophyll-a content of the individual was estimated using anactive chlorophyll fluorometry in a non-destructive manner. After the gametes release, chambers precipitated during culture were dissected and δ13C
and δ18O were analyzed. Data are partly from Takagi et al. (2016a). The decrease of δ13C was accompanied by a decrease in (or almost zero)
photosynthesis at the final stage, indicating that the lack of photosynthesis could be clearly depicted from the chamber-by-chamber isotopic profiles

comparing the δ13C/size gradient at different time slices
with different climates, attention must be paid to the balance of photosynthesis and respiration under each SST. In
any case, to our best knowledge, no “bleaching”
phenomenon due to high temperature has been reported
for modern planktic foraminifera under laboratory culture.
Interestingly, through the macroevolutionary history of
Cenozoic planktic foraminifera, the proportion of photosymbiotic species to the total number of species corresponds well to the pattern of Earth’s climate change(Fig. 17). This proportion became larger during hyperthermal events like the Paleocene/Eocene thermal maximum,
Early Eocene climatic optimum, and MECO. This suggeststhat the photosymbiotic association with algae would makeforaminifera more adaptive to such hot climates and accelerated their speciation. In addition to such adaptivity in theevolutionary scale, it is important to understand theindividual-scale mechanism of photosymbiosis. Since an individual holobiont is involved with photosynthesis,

respiration, and calcification, the net fitness of the photosymbiotic system is difficult to evaluate.
Perturbation in global carbon cycle at the Paleocene/Eocene
boundary

The P/E transition event at 55.5 Ma is often referred to as
the Paleocene/Eocene thermal maximum, characterized
by a huge concomitant perturbation in the global carbon
cycle (Zachos et al. 2001). Large δ13C anomalies observed
in deep-sea sediments, terrestrial soil carbonates, andmammal teeth (associated with a major turnover in landmammal assemblages), was − 3‰ (up to − 6‰) over aninterval of ~ 100–200 kyr near the end of the Paleocene(Kennett and Stott 1991; Norris and Röhl 1999; Pagani
et al. 2006). The duration of these events was estimated to
be no more than 10 kyr, with P/E transition conditions
with a total duration of about 215 ± 5 kyr for δ13C excursion (Röhl et al. 2000) (Fig. 18). Benthic foraminiferal δ18O
values indicate that oceanic deep waters warmed

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

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abFig. 17 Cenozoic climate changes and paleodiversity dynamics of planktic foraminifera. a Cenozoic climate change reconstruction from thestable oxygen isotope profile of deep sea benthic foraminifera (Zachos et al. 2008). b Changes in species number and proportion ofphotosymbiotic species (original data and categorization of photosymbiotic group are from Aze et al. 2011). PETM Paleocene Eocene thermal
maximum, ETM2 Eocene thermal maximum 2, EECO Early Eocene climatic optimum, MECO Middle Eocene climatic optimum, EOT Eocene
Oligocene transition, MMCO Middle Miocene climatic optimum, Paleo Paleocene, Oligo Oligocene, P Pliocene, Q Quaternary. After the rapid
emergence of photosymbiotic species in the Paleocene, the proportion of photosymbiotic species fluctuate with a good correspondence to theclimate. The proportion increased during warmer periods like PETM, EECO, and MECO, which indicate the advantage of photosymbiotic ecology
during such a hot climate

substantially worldwide, from ~ 10 to ~ 18 °C in high latitudes (Kennett and Stott 1991) and from ~ 10 to ~ 16 °C in
the tropics (Bralower et al. 1997). SSTs, as recorded byplanktic foraminiferal δ18O and/or Mg/Ca values, showed
little change (from ~ 26 to ~ 28 °C) at low latitudes but larger increases at high latitudes, leading to diminishedvertical and latitudinal thermal gradients (e.g., Kennett andStott 1991; Kelly et al. 1996).
Actually, the declines in coral reefs could be attributed
to an SST in the tropics that was beyond the maximum
temperature range of corals (Scheibner and Speijer 2008).
In terrestrial environments, a widespread increase inKawahata et al. Progress in Earth and Planetary Science(2019) 6:5

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Fig. 18 Effect of methane release with a δ13C of − 60‰ over 10 kyr on the δ13C value of present-day preindustrial carbon reservoirs. a Responseof δ13C in deep water of the Atlantic, Indian, and Pacific Oceans. Initial values are 0.996‰, 0.620‰, and − 0.508‰, respectively. b Response of δ13C
in warm surface water, cold surface water, and the atmosphere (shifted by + 8‰). Initial values are 2.577‰, 1.355‰, and − 6.367‰, respectively(modified from Dickens et al. 2017)kaolinite was interpreted to herald a warm and wet climate. In association with these environmental changes,
ocean circulation, primary productivity, carbonate dissolution, and the hydrologic cycle showed severely disturbed
wide-ranging environments (Ernst et al. 2006; Stassen

et al. 2012). In order to explain the large negative shift inδ13C value, the dissociation of a large amount of methane
hydrate (with a δ13C of − 60‰) is considered to be
the most plausible mechanism (e.g., Dickens et al.
1997). The amount of carbon required to explain the(2019) 6:5

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negative δ 13C shift is estimated to be as large as2000 PgC (1 PgC = 1015 g of carbon) (Dickens et al. 1997;
Zachos et al. 2005).
In oxygenated condition, the injected methane is easily
converted to CO2 within a few years. Carbon dioxide is
weakly acidic; therefore, it would lead to a reduction in
deep-sea pH, possibly triggering a rapid (< 10 kyr) shoalingof the global CCD by > 2 km, followed by gradual recovery(> 100 kyr) (Zachos et al. 2005). The δ 13C anomaly also
coincides with the most dramatic extinction of 35–50% ofcosmopolitan benthic foraminifera in the past 100 million
years and an apparent change in the dominance of calcareous to agglutinated benthic fauna at some deep seafloorsites (Kawahata et al. 2015). Carbonate saturation level was
more reduced by the higher pressure in the deep-sea.
Such characteristics make the P/E transition event a
possible analog of the modern human perturbation to theocean–atmosphere carbon system. At the P/E transition,
methane was rapidly released to the ocean and atmosphere at an average rate of ~ 0.2 PgC year−1 of CO2 over
10 kyr (Dickens et al. 1997; Zachos et al. 2005). This rate
is approximately 1/30 of the current global emission rate(~ 7.8 PgC year−1) of anthropogenic carbon input.

studies for the Ayeyarwady River were largely overestimated, and they recalculated the CO2 consumption rates
by silicate and carbonate weathering in the entire Himalayan river basin, reporting values of 0.012 PgC year− 1
and 0.024 PgC year−1, respectively. In contrast, France–
Lanord and Derry (1997) proposed that erosion of organiccarbon and its burial in the Bengal Fan may have controlled the global carbon cycle and climate change in thisera.
The estimated global CO2 consumption rates by silicate
and carbonate weathering are 0.13 PgC year−1 and 0.15
PgC year−1, respectively, based on the chemical compositions of major world rivers given by Gaillardet et al. (1999)
and modified data for the Ayeyarwady Rivers by Manaka
et al. (2015a, b). Hartmann et al. (2009) highlighted extremely high silicate weathering rates in small areas suchas the Pacific islands (so-called hyperactive areas and hot
spots), based on a high-resolution global lithological map
and the calculated weathering rate of each rock types.
Current global warming will continue to be a pressing
problem over timescales of decades to centuries because
the anthropogenic CO2 release rate from fossil fuel use
and cement production is 7.8 PgC year−1 (IPCC 2013a),much larger than the CO2 consumption rate of chemical
weathering (Table 2). Anthropogenic land use changes
such as deforestation also lead to a carbon release of 1.1
PgC year−1 to the atmosphere. Currently, a half ofanthropogenic CO2 release remains in the atmosphere(average atmospheric increase: 4 PgC year−1), while approximately a quarter is absorbed into the terrestrial biosphere and the ocean. Terrestrial photosynthesis absorbs
123.0 PgC year−1 CO2, most of which (118.7 PgC year−1)
is returned to the atmosphere via the respiration process.
Net CO2 uptake by the terrestrial biosphere is only 4.3
PgC year−1.
The pCO2 in the global terrestrial waters of lower alluvial basins with well-developed soils and active agriculture is often over several thousand μatm, muchhigher than the atmospheric level (400 μatm) (Cole et al.
2007; Raymond et al. 2013; Manaka et al. 2015b). Terrestrial waters can receive 1.9 PgC year−1 of soil carbon,
which was originally fixed by the photosynthesis
process (Regnier et al. 2013). The global riverine flux ofcarbon to the ocean is 1.06 PgC year−1 (Li et al. 2017),
which could enhance pCO2 in coastal regions.
Soil-derived carbon, mainly dissolved organic carbon, is
decomposed in the water and released directly to theatmosphere in the form of CO2 (Rasilo et al. 2017). The
global CO2 degassing rate from water is 1.1 PgC year−1,
which is comparable in magnitude to that from fossil
fuel and cement. One of the largest CO2 evasions is reported for the Amazon River; the estimated flux is ashigh as 0.47 PgC year−1 from the entire basin (Richey
et al. 2002).

Kawahata et al. Progress in Earth and Planetary Science

Long-term neutralizing process against enhanced pCO2

Terrestrial waters stimulate physical and chemical
weathering of both silicate and carbonate rocks andtransport HCO3− as well as other forms of dissolved andparticulate carbon from land to the ocean (Eqs. 2 and 3)(e.g., Hecht and Oguchi 2017). In particular, silicate
chemical weathering is one of the most important processes working as a net sink of CO2. Therefore, it alsofunctions as neutralizing process against OA.
It is generally assumed that one of the most active
chemical weathering scenarios globally occurs in the highHimalaya, as a result of the ongoing uplift of the Himalayan–Tibetan Plateau and the heavy rains associated with
the Indian–Asian monsoon, which is also linked with theNorth Pacific monsoon (e.g., Tada and Murray 2016; Tada
et al. 2016; Clift 2017; Dado and Takahashi 2017).
Gaillardet et al. (1999) compiled dissolved major ion data
in Himalayan Rivers and reported high chemical weathering rates for silicate in the Ganges–Brahmaputra andAyeyarwady Rivers, and for carbonate in the Yangtze
River. Raymo et al. (1988) and Raymo (1991) focused on
the Sr isotope record in marine sedimentary cores andfound that active silicate weathering in this area may havetriggered global cooling during the Eocene. However, theimportance of a long-term CO2 sink by silicate weatheringin this area has been questioned in recent studies. Bickle
et al. (2001, 2003) reported that Sr isotopes cannot be
used as a proxy for silicate weathering owing to regional
metamorphism in this area. Manaka et al. (2015a) notedthat the silicate weathering rate reported in previous

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

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Table 2 Global carbon exchange between land, inland waters, ocean, and atmosphere since the industrial era
Pg C year−1

References

CO2 Consumption by chemical weathering of silicates in Himalayan River Basins

0.012

Manaka et al. (2015a, b)

CO2 Consumption by chemical weathering of carbonates in Himalayan River Basins

0.024

Global CO2 consumption by chemical weathering of silicates

0.132

Global CO2 consumption by chemical weathering of carbonates

0.149

CO2 release to the atmosphere by fossil fuel use and cement production

7.8

CO2 release to the atmosphere by anthropogenic land Use change

1.1

Total terrestrial photosynthesis

123.0

Total terrestrial respiration

118.7

Gaillardet et al. (1999); Manaka et al. (2015a, b)

IPCC (2013)

Net atmospheric CO2 uptake by terrestrial biosphere

4.3

Total soil carbon input from land to inland waters

1.9

Regnier et al. (2013)

Global riverine flux of carbon to the ocean

1.06

Li et al. (2017)

CO2 release from inland waters to the atmosphere1.1

Regnier et al. (2013)

For future environments, global warming is expected
to enhance both chemical weathering (IPCC 2013a) andsoil respiration (Hashimoto et al. 2015). This warming ispredicted to cause extreme rainfall events, particularly in
the Himalayan lowland alluvium, which may further increase pCO2 of river water (Goswami et al. 2006).
Anthropogenic eutrophication in aquatic ecosystems isanother widespread problem globally (Mackenzie et al.
2002). Cai et al. (2011) reported that large riverine inputs of nutrients, as well as OM and dissolved CO2, canenhance the acidification of subsurface coastal waters. In
contrast, this eutrophication can enhance photosynthesis, and thus CO2 uptake, in enclosed conditions suchas lakes, which may be important to the carbon cycle,
particularly in developing countries (Manaka et al.
2013). Therefore, changes in water composition in response to environmental change should be borne inmind in the future.
Inventory of alkalinity required for the preservation of
deep-sea carbonate

In order to understand the relationship between carbonate dissolution and pCO2 in deep sea, a simple ocean
carbon cycle model was used based on the three-box
model of Toggweiler and Sarmiento (1985). The details
are described in Yamamura et al. (2007) (Fig. 19). The
model has five state variables: phosphate, oxygen, potential temperature, alkalinity, and DIC. Salinity is fixed to
35 practical salinity units (psu), to which phosphate, alkalinity, and DIC are normalized. Gas exchanges of CO2
and oxygen with the atmosphere are driven by the respective partial pressure gradients across the air–sea
interface and a “piston velocity” of 3 m day−1 (Broecker
and Peng 1982). The carbonate saturation depth at theintersection of [CO32−] in the deep box and [CO32−]sat(the carbonate ion concentration that is at saturation

with respect to mineral calcite) is calculated from theapparent solubility product Ksp for calcite. The production rate of organic carbon from the surface ocean isexpressed as the flux of phosphate, which is assumed tooccur with a stoichiometry of P:Corg:-O2 = 1:117:170(Anderson and Sarmiento 1994). With the production
rate, the global average export ratio is adjusted to
CaCO3:Corg = 0.1 based on Sarmiento et al. (2002).
For initial values, which represent modern conditions,the model predicts [CO32−] of about 94 mol kg−1 and 2 °C
in the deep box (Figs. 19 and 3). The calcium carbonate
saturation horizon under these model conditions is located at a depth of about 3700 m. The mid-Cretaceous isbelieved to be 6–12 °C warmer than today and had 2–10
times more atmospheric pCO2 (e.g., Berner and Kothavala2001). The model results also indicate that carbonate dissolution in the northwestern Pacific during the Albian
was severe (Fig. 3). However, this is contrary to the geological observation that carbonate was apparently
present at paleodepths of ~ 1000–2000 m in DSDP/
ODP cores (e.g. Dean 1981). This apparent contradiction may be resolved if there were more alkalinity inthe oceans during the Albian, mid-Cretaceous, than
today. Global warming with high pCO2, like during the
Albian, would have induced intensified hydrologicalcycles on the Earth’s surface (e.g., Jenkyns 2010). Subsequently, this would have resulted in intensified weathering and increased the supply of nutrients and alkalinity(Eqs. 2 and 3) (Fig. 2).
To investigate the effect of a higher alkalinity content
in the ocean, a set of model experiments was conductedwith the same range of high latitude surface box temperatures and four-times the present atmospheric pCO2
(Fig. 19). In the late Albian, the deep-sea temperature
was estimated to have been ~ 17 °C in the equatorial Pacific (Huber et al. 2002) and pCO2 was suggested to varyKawahata et al. Progress in Earth and Planetary Science(2019) 6:5

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Fig. 19 Three-box model sensitivity of saturation depth and carbonate ion concentration at deep box at a high CO2 concentration (1020 ppm)
(Yamamura et al. 2007). Approximately 15% of the world ocean surface area is represented by the high-latitude box. Variable T represents thelarge scale meridional overturning of the world ocean. Polar vertical mixing is meant to represent deep water ventilation and formation aroundthe Antarctic continent today. Phosphate serves as the main nutrient accounting for organic carbon production. The phosphate concentration in
the low-latitude surface box drives its new production by restoring the phosphate concentration to its initial value. The production rate of organiccarbon from the surface ocean is expressed as the flux of phosphate, which is assumed to occur with a stoichiometry of P: Corg:-O2 = 1:117:170(Anderson and Sarmiento 1994). In this model, the new production rate at the low-latitude surface box, which is described as Pl, depends on the
upwelling supply flux of phosphate due to the thermohaline circulation. In the high-latitude box, the new production rate is fixed to a prescribed value
since the productivity seems to be affected by factors other than nutrients. With the production rate, we adjust the global average export ratio of
CaCO3/Corg to 0.1 based on Sarmiento et al. (2002). The solid line indicates the carbonate ion concentration (μmol/kg) and the broken line indicatessaturation depth (m). The light-gray shading indicates the saturation depth below 0 m. The dark shading indicates the range of the saturation depth,
similar to the interglacial value

between approximately two and five times more thanpresent-day atmospheric pCO2 (Bice et al. 2006). In sucha case, the required alkalinity inventory to maintain thelevel of saturation depth (deep than 2000 m) at theinterglacial condition is greater than 1.2 times (Fig. 19).
In any case, the model results indicate that under highpCO2 during the mid-Cretaceous, an increase in theoceanic inventory of alkalinity is likely necessary to preserve any carbonate on the seafloor (Fig. 2; Table 1).
Otherwise, the entire ocean may be undersaturated withrespect to CaCO3.

Key factors inducing ocean acidification

It is commonly believed that increasing pCO2 will induce an enhanced greenhouse effect and change seawater chemistry, resulting in future OA (Figs. 2 and 3;
Table 1) (IPCC 2017). The mid-Cretaceous represents
one of the warmest climate intervals during the entire
Phanerozoic (termed the “hot house”) due to enhanced
pCO2 (e.g., Takashima et al. 2009). Actually, themid-Cretaceous is believed to have been a much warmer
environment with much higher pCO2 (e.g., Berner andKothavala 2001). The fact that we see relatively better

Kawahata et al. Progress in Earth and Planetary Science(2019) 6:5

Modern & P/E boundary

rapid ocean acidification

Page 30 of 37

Cretaceous

neutralization bychemical weathering

Fig. 20 Schematic diagram of ocean acidification and the buffering system in the Earth’s surface system. Chemical weathering is important andacts as a neutralizing process. The level of CO2 is a secondarily important factor. The most important factor is “Changing SPEED.” A natural neutralizing
system such as chemical weathering cannot catch up too rapid a global environmental change under modern conditions or at the P/E boundary(Kawahata et al. 2015)

carbonate preservation in the mid-Cretaceous suggeststhat the ocean surface was sufficiently basic, despite
higher pCO2, to allow the growth of calcifying producers. Alkalinity should have been supplied by chemical
weathering on the land. A higher alkalinity inventory inthe ocean is in fact expected according to an enhancement of the chemical weathering of silicate rocks, driven
by pCO2 and intensified hydrological cycles on the
Earth’s surface (Fig. 2) (e.g., Broecker and Sanyal 1998;
Cohen et al. 2004).
The biogeochemical cycle on the Earth’s surface yields
a buffering system for pH and carbonate ion change(Fig. 20). Therefore “high speed” and “high level” pCO2
are the primary and secondary factors, respectively, leading to acidification in the Earth surface’s system. Too
rapid a release of anthropogenic CO2 (today and at P/Eboundary) will yield severe OA in deep ocean. In combination with the high solubility of carbonate minerals(Fig. 1), benthic fauna will thus be threatened.
Concluding remarks and future investigation

In the global carbonate budget, foraminifera and corals
are two major calcifiers contributing to the burial of carbonate in sediments/reefs. Our extensive review of thesecalcifiers suggests similar, but slightly different, responsesof two representative species to ocean warming and acidification. Both calcifiers generally show negative responses
to OA (increasing pCO2 and decreasing saturation state),
although their responses are dependent upon life stage
and taxa. The response of corals to heat stress is variable
among species. The response of corals to acidified

seawater may depend on the degree of seawater transport
into the site of calcification (calcifying fluid), but this interpretation remains controversial. Foraminifera, which
actively pump protons out from the site of calcification
and are surrounded by a low external pH, may deal withcalcification under elevated pCO2 using an energy cost
approach.
Ocean warming (i.e., higher temperature) likely induces the bleaching of symbiont-bearing marine calcifiers, including corals and foraminifera, and results inthe suppression of calcification of both groups; death of
the host organisms may result. However, symbiotic foraminifera appear more robust and resilient than symbiotic corals at higher temperatures. During warm periodsof the Earth’s history, when SSTs and deep-water temperatures exceeded those of modern oceans, foraminifera survived under extremely warm, high pCO2, anddysoxic conditions. However, corals and living reefs
never existed during these warm periods, known as thereef gaps (Veron 2008). Therefore, benthic foraminifera
seemed more adaptive and robust than corals to ocean
warming and acidification over the Earth’s history; they
will likely remain so in the future. On the other hand,during a possible OA event at 65 Ma at the end
Cretaceous mass extinction, all remaining ammonites
and belemnites became extinct, as did a high percentage
of bivalves, gastropods, and echinoids, and almost allplanktic foraminifera. Scleractinian corals appear to havejust barely survived and 1 (the Faviidae) retained 6 of its
original 16 genera; the others survived with only 1 or 2.
Therefore, in order to pursue the issue of adaption and(2019) 6:5

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robustness, it is required to further understand themechanisms underlying variations in sensitivity to heatstress and acidified seawater for both corals andforaminifera.
Since the pCO2 in the atmosphere is currently increasing, pCO2 and DIC in seawater are increasing, while pH
and the saturation state of carbonate minerals are decreasing without any change in total alkalinity. However,
over the long-term, it is well known that alkalinity has
fluctuated largely in the Earth’s surface environment.
Therefore, it is important to quantitatively reconstructalkalinity, which is another key factor determining thesaturation state of carbonate minerals.

extracellular pH distribution shows that proton
pumping produces a high internal pH and large
internal–external pH gap associated with
foraminiferal calcification.
6. X-ray micro-computed tomography, which canestimate shell density, indicates that the dissolutionof Globigerina bulloides shells begins at the central
area of the shell, whereas the outer calcite layer,with a euhedral crystalline structure, is resistant
to dissolution.
7. Alkalinity, another key factor controlling OA, has amuch longer residence time. The “high speed” and“high level” of pCO2 are the primary and secondary
factors, respectively, leading to severe OA in the
Earth surface’s system. Too rapid a release ofanthropogenic CO2 (today and at P/E boundary)would bring severe OA, even in the deep ocean.

Kawahata et al. Progress in Earth and Planetary Science

Summary
Global warming and OA are often referred to as the “eviltwins” of climate change due to enhanced pCO2. Bothwill yield severe threats to marine calcifiers in the near
future (Fig. 1)
1. CO2 dissolution into seawater increases DIC andpCO2 and decreases pH, with no change to total
alkalinity. While OA may progress ubiquitously, thecorresponding changes in pH and pCO2 vary locally.
Since pCO2 is sensitive to temperature (~ 4% increase
per 1 °C), OA is more critical in the polar and highlatitude regions, as compared with temperate andtropical regions. In addition, older deep-water hasenhanced acidity owing to the addition of CO2 via
the degradation of OM, which promotes carbonate
dissolution via a synergistic effect with high pressure.
2. Anthropogenic CO2-driven OA will pose a severe
threat to marine calcifying organisms because thecalcification rates of organisms generally decrease inacidified seawater. However, some opposite responses ofreef-dwelling larger benthic foraminifera are reported.
3. In the case of the coral calcification response of
Aropora digitifera to acidified seawater, calcification
tended to decrease according to the increase of
pCO2 in both aposymbiotic and symbiotic polyps,
although the calcification of adult branches was not
reduced markedly in higher pCO2 conditions. Thissuggests that OA response may depend upon life
cycle stage.
4. As exemplified by reef-building corals and reef-dwelling
larger benthic foraminifera, photosymbiosis is regarded
as an adaptive ecology for living in warm, oligotrophic,
and well-lit oceans. Bleaching is likely a global threat tomarine symbiont calcifying organisms, especially at hightemperature and enhanced light intensity.
5. Based upon Ca isotope measurements of coralskeletons, the transport path of Ca2+ and the extent
of its activity would predominantly control the
carbonate precipitation rate. Visualization of theAbbreviations
CCD: Carbonate compensation depth; DIC: Dissolved inorganic carbon;
kyr: Kilo years; Ma: Megaannum (one million years); OA: Ocean acidification;
OM: Organic matter; pCO2: Partial pressure of CO2; SST: Sea surface
temperature; Ωarag: Aragonite saturation state; Ωcal: Calcite saturation state
Acknowledgements
This research was partly supported by the Japan Society for the Promotionof Science Grants-in-Aid for Scientific Research (A) to H. Kawahata (15H02139).
Funding
This work was supported by JSPS KAKENHI Grant Numbers 22224009,
15H02139, 15H02813, 18H03366, and 18 K18793.
Availability of data and materials
Please contact author for data requests. All authors will share data and materials.
Authors’ contributions
HK proposed and designed the study. KF described the part of the manuscript,
including the topic on the bleaching in algal symbiont-bearing large benthic
foraminifera. AI described the part of the manuscript, including the topic on thebleaching in corals. MI described the part of the manuscript, including the topic
on biomineralization of corals based on geochemical analyses. SI described thepart of the manuscript, including the topic on deep-sea carbonate dissolution
and progressive foraminiferal shell dissolution. AK described the part of themanuscript, including the topic on Response of foraminifera against ocean
acidification. AM described the figures. TM described the part of the manuscript,
including the topic on long-term neutralizing process against enhanced pCO2.
KM described the part of the manuscript, including the topic on the extreme
warmth in the Cretaceous greenhouse system. HT described the part of themanuscript, including the topic on warmth in the Middle Eocene Climatic
Optimum (MECO) event and foraminiferal photosymbiosis. TT described thepart of the manuscript, including the topic on biomineralization of foraminifera.
TY described the part of the manuscript, including the topic on characteristics
of biotic and abiotic carbonate minerals. AS described the part of themanuscript, including the topic on general feature of marine carbonate system.
All authors interpreted the data, read, and approved the final manuscript.
Authors’ information
No special information for each author.
Competing interests
The authors declare that they have no competing interests.

Publisher’s NoteSpringer Nature remains neutral with regard to jurisdictional claims in publishedmaps and institutional affiliations.

(2019) 6:5

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Author details
1
Atmosphere and Ocean Research Institute, The University of Tokyo, 5-1-5
Kashiwanoha, Kashiwa, Chiba 277-8564, Japan. 2Geological Survey of Japan,
National Institute of Advanced Industrial Science and Technology (AIST),
Higashi 1-1-1, Tsukuba, Ibaraki 305-8567, Japan. 3Department of Physics andEarth Sciences, Faculty of Science and Tropical Biosphere Research Center,
University of the Ryukyus, 1 Senbaru Nishiharacho, Okinawa 903-0213, Japan.
4
Department of Earth and Planetary Science, Faculty of Science, Okayama
University, 3-1-1 Tsushima-naka, Okayama 700-8530, Japan. 5Japan Agency
for Marine-Earth Science and Technology (JAMSTEC), 2-15, Natsushima,
Yokosuka 237-0061, Japan. 6The Center for Academic Resources andArchives, Tohoku University Museum, Tohoku University, 6-3 Aramaki-Aza
Aoba, Aoba-ku, Sendai, Miyagi 980-8578, Japan. 7Department of Forest Soils,
Forestry and Forest Products Research Institute, 1 Matsunosato, Tsukuba,
Ibaraki 305-8687, Japan. 8Department of Earth Sciences, School of Education,
Waseda University, 1-6-1 Nishiwaseda, Shinjuku, Tokyo 169-8050, Japan.

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