Research Paper On Ocean Acidification Graph


Both warm- and cold-water corals secrete calcium carbonate skeletons that build up over time to create a three-dimensional reef matrix that provides habitat for thousands of fish and other species. The production of limestone-like calcium carbonate is high enough in many warm-water coral reefs to establish carbonate structures. High rates of calcification are sufficient to overcome significant rates of bioerosion and wave driven physical erosion. These structures underpin the framework of barrier reefs and islands, which are critically important to tropical coastlines. Although they occupy less than 0.1% of the ocean floor, tropical coral reef ecosystems provide habitat for at least 25% of known marine species, with many reef species still to be discovered (Fisher et al., 2015). The biological diversity of warm-water coral reefs has been estimated to include ~1–9 million species that live in and around coral reefs (Reaka-Kudla, 1997, Census of Marine Life, In deeper parts of these warm-water reef systems, the tendency toward carbonate dominated reef structures diminishes as light levels decrease (Bongaerts et al., 2010a). At low light levels, erosion and dissolution exceed calcium carbonate production, leading to coral communities that may be abundant yet with little or no three-dimensional calcium carbonate reef framework. Extending from 40 to 150 m, these “mesophotic” (low light) coral reefs also provide extensive habitat, with the rates of discovery of species remaining very high due to these reefs being difficult to visit (Bongaerts et al., 2010a, 2011). Mesophotic reef systems probably cover a comparable area to shallow warm-water coral reefs (Bongaerts et al., 2010a; Slattery et al., 2011).

Both shallow or deeper mesophotic coral reefs are dominated by scleractinian corals that form symbiosis with dinoflagellate protists from the genus, Symbiodinium. On the basis of this symbiosis, their intracellular symbionts (i.e., living within the gastrodermal or digestive tissues of their coral hosts) are able to photosynthesize and provide the host coral with a rich source of sugars, glycerol, lipids, and other organic compounds (Muscatine, 1990). This relationship enables corals to grow and calcify at high rates in the clear, warm, and shallow water conditions along tropical coastlines (Muscatine and Porter, 1977). The abundance of Scleractinian corals hosting Symbiodinium decreases with depth beyond 20–40 m, depending on the clarity of the water column. The deepest Scleractinian corals that are symbiotic with Symbiodinium, are found 100 m or more below the surface of tropical waters (Englebert et al., 2014). The productivity of this symbiosis is complemented by the ability of corals to capture and feed on waterborne particles and plankton (i.e., polytrophy). The combined ability to photosynthesise, as well as feed, underpins the success of the highly productive coral reef ecosystems that line many tropical coastlines. Evidence from isotope signatures within fossils reveal that Scleractinian corals have been symbiotic with Symbiodinium for over 230 million years (Stanley and Fautin, 2001; Muscatine et al., 2005), most probably driving productive and diverse ecosystems that were not too different from those of today.

Cold-water coral reefs extend to depths of 3,000 m although some cold-water corals can be found growing in waters as shallow as 50 m (e.g., Norwegian shelf). Below 200 m depth there is so little light that photosynthesis is no longer possible. As a result, cold-water corals do not form a symbiosis with Symbiodinium and depend instead on particle feeding. Discoveries of the locations and extent of cold-water reefs has primarily been driven by advances in underwater technologies for surveying and mapping (Turley et al., 2007; Ramirez-Llodra et al., 2010). For example, vast extents (~2,000 km2) of cold-water coral reefs, some shown to be thousands of years old (>8,000 years), have been found in Norwegian waters in past decades (Fosså et al., 2005). Cold-water coral reefs have now been discovered in every ocean, forming important assemblages within the deep ocean that provide critical habitat to thousands of other species, including many commercially important species.

Human communities derive many benefits from coral reefs including food, income, recreation, coastal protection, cultural settings, and many other ecological goods and services (Cinner et al., 2009; Costanza et al., 2014). Despite their biological diversity, productivity and importance to humans, both warm and cold-water coral reefs are being heavily impacted by human activities due to both local and global influences (Hall-Spencer et al., 2002; Burke et al., 2011). As a result, many coral reefs are rapidly declining across the world. While local factors can have significant impact on coral reefs (e.g., pollution, overfishing, and the physical destruction of reefs), changes in ocean temperature and chemistry due to anthropogenic activities are dramatically reducing the distribution, abundance, and survival of entire coral reef ecosystems (Gattuso et al., 2014b; Hoegh-Guldberg et al., 2014). Given these risks and the importance of coral reefs to humans and marine biodiversity, the present paper focuses on the challenges that warm and cold-water coral reef ecosystems and their human communities are facing, particularly those posed by rapidly warming and acidifying oceans.

Distribution, Abundance, and Importance of Coral Reef Ecosystems

Warm-water coral reefs are prominent ecosystems within coastal areas of the Pacific, Indian, and Atlantic oceans (Figures 1A,B), where they are typically found in a broad band (30°S to 30°N) of warm, sunlit, alkaline, clear, and relatively nutrient deficient ocean waters (Kleypas et al., 1999b). Here, Scleractinian or reef-building corals proliferate, depositing copious amounts of calcium carbonate. As corals die, their dead skeletons build up over time and are “glued” together by the activities of other organisms such as encrusting red coralline algae (Glynn and Manzello, 2015). Other organisms such as calcifying green algae, invertebrates, and phytoplankton also contribute to the overall carbonate budget of warm water coral reefs (Hutchings and Hoegh-Guldberg, 2009), leading to three-dimensional calcium carbonate structures that build up over hundreds and thousands of years. In turn, the three-dimensional structures (Figure 1C) within warm-water reef systems creates habitat for hundreds of thousands of species, many of which support coastal human populations with food, income, and other ecological goods and services such as coastal protection. Coral reefs are also important sources for bio-prospecting and the development of novel pharmaceuticals. The asset value of coral reefs has been estimated as close to $1 trillion (Hoegh-Guldberg, 2015) with the economic value of goods and services from coral reefs exceeding $375 billion annually, with benefits flowing to over 500 million people in at least 90 countries worldwide (Burke et al., 2011; Gattuso et al., 2014b).

Figure 1. (A) Distribution of warm-water and cold-water coral reefs (credit: Hugo Ahlenius, 2008, UNEP/GRID-Arendal, (B) Location of warm-water coral reef cells and provinces, from Hoegh-Guldberg et al. (2014). (C) Warm-water carbonate coral reef from the Great Barrier Reef, Australia (credit: Ove Hoegh-Guldberg). (D) Mesophotic coral community of North Sulawesi, Indonesia. (Credit: Pim Bongaerts, University of Queensland). (E) Deep-water community of Lophelia pertusa from the Mississippi Canyon at ~450 m depth (Image from NOAA, licensed under the Creative Commons Attribution-Share Alike 2.0 Generic license).

As light levels decrease with depth, decalcification dominates and the overall carbonate balance of reef ecosystems shifts to negative (Barnes and Chalker, 1990; Bongaerts et al., 2010a). Under these conditions, Scleractinian corals and their symbionts persist with reefs being referred to as “mesophotic” (Bongaerts et al., 2010a, 2011; Robinson C. et al., 2010). In these habitats, colonies of Scleractinian corals are often platelike in shape, orientating themselves to maximize light harvesting under these dim light conditions (Figure 1D). Mesophotic reef systems are also primarily restricted to areas where water clarity, carbonate ion concentrations, and temperatures are relatively high. Like their counterparts in shallower regions, mesophotic reef systems play an important role in supporting fisheries and hence human livelihoods. Given the difficulty of working at depths of more than 30 m (beyond SCUBA-diving depth) many species remain to be discovered (Bongaerts et al., 2010a). Mesophotic reefs therefore have an unknown potential to be sources of novel pharmaceuticals and other potentially beneficial compounds (Leal et al., 2012). As a result, their true value has probably been underestimated.

Cold-water corals generally form reefs at much greater depths from 200 to 2,000 m however in some regions they are found at shallower depths (Fosså et al., 2002; Freiwald et al., 2004). Deep-water corals are not dependent on light levels as they are not symbiotic with Symbiodinium. Due to the colder and more CO2 rich (and hence less alkaline) waters, deep-water corals grow slower than warm-water corals, forming aggregations that are variously termed patches, banks, thickets, bioherms, mounds, gardens, and massifs. In the absence of significant wave action, these fragile and slow growing reefs form aggregations that can cover vast tracks of the seabed (e.g., 2,000 km2 in Norwegian waters (Hall-Spencer et al., 2002) and involve near mono-specific stands of Scleractinian corals such as Lophelia pertusa and Oculina varicosa (Figure 1E). In addition to Scleractinian corals, they often exhibit a wide variety of abundant coral-like organisms, including soft corals, gorgonians, and Alcyonaceans.

Recent Changes in the Extent of Anthropogenic Stresses on Coral Reef Ecosystems

Coral reefs are facing growing challenges from the local to global effects of human activities. Over the past 200 years, human activities have fundamentally changed coastlines, overexploited resources such as fish stocks, and polluted coastal waters, to a point where many coral reef ecosystems are degrading rapidly (Jackson et al., 2001; Pandolfi et al., 2003; Hoegh-Guldberg, 2014b). Warm-water coral reefs, for example, have declined by at least 50% over the past 30–50 years in large parts of the world's tropical regions (Hughes, 1994; Gardner et al., 2003; Bruno and Selig, 2007; De'ath et al., 2012). Similar conclusions have been reached for cold-water reefs where human activities have put these systems under escalating pressure from the mid-1980s onwards. Key drivers of the destruction of cold-water reefs include commercial bottom trawling, hydrocarbon exploration and production, deep sea mining, cable and pipeline placement, pollution, waste disposal, coral exploitation, and trade, and destructive scientific sampling (Hall-Spencer et al., 2002; Turley et al., 2007; Roberts and Cairns, 2014). The increase in impacts from human activities is a result of rapid advances in technologies for visualizing and exploiting the biological and mineral resources of deep water habitats (Freiwald et al., 2004; Ramirez-Llodra et al., 2010). Many populations of deep-sea corals (Scleractinians, gorgonians) have very slow turn-over rates and may live for centuries, with some species such as black corals (Antipatharians) living for thousands of years. The longevity and slow growth rates of these taxa means that recovery from anthropogenic stressors will be very slow. The areas inhabited by the deep-sea reefs are also a “resource frontier” for hydrocarbon extraction and mining of high value and “high-tech” metals (Roberts and Cairns, 2014). Hence, it is likely that anthropogenic impacts on these reefs will expand. These impacts are also likely to interact with ocean warming and acidification (Figure 2A), which pose growing and serious risks to coral reef ecosystems on their own. The direct impact of these changes to coral reefs have been growing since the early 1980s (Hoegh-Guldberg et al., 2007, 2014; Eakin C. M. et al., 2010; Gattuso et al., 2014b). The latter are the direct result of the burning of fossil fuels and have been driving growing impacts on warm water coral reefs since the early 1980s. Understanding and solving both local and global threats to coral reefs will be critically important if they are to survive some of the greatest rates of environmental change in the recent history of the Earth (Hönisch et al., 2012; Pörtner et al., 2014).

Figure 2. (A) Linkages between the build-up of atmospheric CO2 and the slowing of coral calcification due to ocean acidification. Approximately 30% of the atmospheric CO2 emitted by humans has been taken up by the ocean (IPCC, 2013) where it combined with water to produce carbonic acid, which releases a proton that combines with a carbonate ion. This decreases the concentration of carbonate, making it unavailable to marine calcifiers such as corals. (B) Temperature, [CO2]atm, and carbonate-ion concentrations reconstructed for the past 420,000 years. Carbonate concentrations were calculated (Lewis et al., 1998) from [CO2]atm and temperature deviations from conditions in the decade of the 2000s with the Vostok Ice Core data set (Petit et al., 1999), assuming constant salinity (34 parts per trillion), mean sea temperature (25°C), and total alkalinity (2,300 mmol kg−1). Acidity of the ocean varies by ± 0.1 pH units over the past 420,000 years (individual values not shown). The thresholds for major changes to coral communities are indicated for thermal stress (+2°C) and carbonate-ion concentrations ([carbonate] = 200 μmol kg−1, approximate aragonite saturation ~Ωaragonite = 3.3; [CO2]atm = 480 ppm). Coral Reef Scenarios CRS-A, CRS-B, and CRS-C are indicated as A, B, and C, respectively, with analogs from extant reefs. Red arrows pointing progressively toward the right-hand top square indicate the pathway that is being followed toward [CO2]atm of more than 500 ppm. From Hoegh-Guldberg et al. (2007) with permission of Science Magazine.

Warm-water coral reefs are largely dependent on the physical and chemical changes occurring in the surface of the ocean, whereas cold-water reef systems are tied relatively more to the broad scale conditions of the bulk ocean (Freiwald et al., 2004; Eakin C. M. et al., 2010). In this respect, there are likely to be differences in terms of the rate and characteristics of the changes that are occurring. These differences also translate into different trajectories when it comes to near and long-term projections of planetary warming and ocean acidification.

Warm-water coral reef environments have experienced relatively small amounts of variability in terms of temperature and carbonate ion concentrations, even with the relatively substantial swings in average global temperature and atmospheric CO2 concentration during the glacial cycle (Figure 2B). Warm-water coral reefs contracted toward the equator during glacial periods, and re-expanded along the tropical and subtropical coastlines of the world during the intervening warm periods (Hubbard, 2015). While these changes were rapid relative to geological time frames, they occurred over periods of 10,000 years or more and are slow when compared to climatic changes that have occurred since pre-industrial. While our understanding of how conditions have changed in terms of the habitat of deep-water coral reefs over geological time is limited, it is very likely that conditions varied even less over these long periods than those surrounding the warm-water coral reefs.

It is virtually certain that the upper ocean has warmed between 1971 and 2010 and likely that it has warmed between 1870s and 1971 (IPCC, 2013). These changes are consistent with those expected from the associated rise in greenhouse gas concentrations in the atmosphere (IPCC, 2013). The average sea surface temperatures (SST) of the Indian, Atlantic, and Pacific oceans have increased by 0.65, 0.41, and 0.31°C during 1950–2009 (Table 30-1 in Hoegh-Guldberg et al., 2014). The influence of long-term patterns of climate variability such as the Pacific Decadal Oscillation (PDO) contribute to variability at regional scales and confound efforts to detect and attribute regional changes to anthropogenic greenhouse gas emissions (Hoegh-Guldberg et al., 2014). Nonetheless, examination of the Hadley Centre HadISST1.1 data (Rayner et al., 2003) over 60 years (1950–2009) reveals significant warming trends in SST for many sub-regions of the ocean (Table 30-1 in Hoegh-Guldberg et al., 2014). Significant trends are clearly demonstrated within the six major warm-water coral reef regions, with the exception of the Gulf of Mexico/Caribbean Sea region (Table 1). Rates of increase in SST in the warm-water coral reef regions range from 0.07°C (west Pacific Ocean) to 0.13°C (Coral Triangle and southeast Asia) per decade, resulting in an overall increase in the regions of between 0.44 and 0.79°C during the period from 1950 to 2009.

Table 1. Changes in sea surface temperature (SST) in six major warm-water coral reef provinces (Figure 1B) over the period 1950–2009 using 1 × 1 degree monthly SST data extracted from the Hadley Centre HadISST1.1 data set (Rayner et al., 2003).

In addition to the heat content and temperature of the upper layers of the ocean, the research community is virtually certain that ocean chemistry is also changing as a result of the increasing amounts of CO2 entering the Ocean (IPCC, 2013). Observed increases in salinity at tropical latitudes are consistent with the amplification of the global hydrological cycle (Durack and Wijffels, 2010; Durack et al., 2012), including rainfall, which have significant implications for coastal ecosystems such as warm-water coral reefs. At regional levels, changes in storm and rainfall intensity also have the potential to influence coastal water quality, which is important to coral reefs, as a result of the interplay between droughts, coastal and catchment erosion, and sudden inundation (flood) events. The impact of climate change adds to those from other human activities that are already impacting water quality, coastal erosion and biological systems.

Average global sea levels are increasing by an average of 3.2 mm year−1 (over 1993–2010) as a result of warming of the ocean (thus increasing volume) and the melting of land ice (IPCC, 2013). Sea level rise varies between regions as a result of differences in local oceanography and geology and the influence of long-term variation in regional climate. Some areas that have significant warm-water coral reefs, such as Southeast Asia and northern Australia, have reported rates of sea level rise of around 10 mm year−1. While the direct attribution of changes in regional wind strength, storm intensity and frequency to global warming is challenging due to long-term variability, there is considerable evidence that the frequency and intensity of the strongest tropical storms in some regions (e.g., North Atlantic; IPCC, 2013) has increased since the 1970s. The combination of higher sea levels and more intense storm systems is likely to increase the amount of force exerted by wave action on coastal areas, which has implications for coastal infrastructure, as well as the state of ecosystems such as coral reefs, mangroves, and seagrass beds (Hamylton et al., 2013; Saunders et al., 2014).

Changes have also occurred in the pH of ocean surface waters over the past 100 years, a phenomenon which is referred to as ocean acidification (Kleypas et al., 1999a; Caldeira and Wickett, 2003; Gattuso et al., 2014a). As CO2 enters the ocean, it reacts with water increasing hydrogen ion concentration (thus decreasing ocean pH) and decreasing the carbonate ion concentration. While the overall change in ocean pH appears small (0.1 pH units over the past 150 years), this is actually a 26% increase in the concentration of hydrogen ions. Experimental evidence shows a reduction in carbonate ions with ocean acidification is biologically significant, since it can affect the rate at which marine organisms, such as corals build their calcareous structures (Kroeker et al., 2013). However, understanding of the mechanisms driving the sensitivity of coral calcification to ocean chemistry, such as the response of the pH of the internal calcifying fluid in which the coral skeleton forms to the concentration of dissolved organic carbon, are only being untangled (Comeau et al., 2017). These changes in ocean chemistry are temperature dependent, with the CO2 absorption and consequently acidification being highest when waters are cooler. The aragonite (one form of calcium carbonate) saturation state (Ωarag) is essentially the ratio between the concentrations of calcium and carbonate ions (Doney et al., 2009). The aragonite saturation state shows a similar distribution to sea surface temperature with Ωarag being highest in the warmest ocean regions and lowest in polar regions (Jiang et al., 2015). Surface waters of the ocean are generally supersaturated with respect to aragonite (Ωarag > 1). However, in warmer waters where Ωarag is not projected to fall to <1 (thus undersaturated with respect to aragonite, Figure 3), substantial impacts are likely to still occur on calcifying organisms. There is substantial evidence that carbonate accretion on warm-water coral reefs approaches zero or becomes negative when Ωarag falls below 3.3 (Hoegh-Guldberg et al., 2007; Chan and Connolly, 2013), a level likely to be reached in tropical surface waters within the next few decades at current rates of greenhouse gas emission (Hoegh-Guldberg et al., 2007; Ricke et al., 2013).

Figure 3. Aragonite saturation state of the surface ocean simulated by the University of Victoria Earth System Model under different atmospheric concentrations of CO2. 280 ppm represents pre-industrial and 394 ppm levels in 2012. Four hundred and fifty ppm is projected to be reached during 2030s under Representative Concentration Pathway (RCP) 4.5, 6.0, and 8.5, and to approach, but not reach 450 ppm, during 2040s under RCP 2.6 (IPCC 2013). Eight hundred ppm is projected to be reached during 2080s under RCP 8.5 only. Fields are calculated from the model output of dissolved inorganic carbon concentration, alkalinity concentration, temperature, and salinity, together with the chemistry routine from the OCMIP-3 project. Modified from Figure SM30-2 in Hoegh-Guldberg et al. (2014; reprinted with permission of IPCC AR5).

The global distribution of cold-water corals is at least partly limited by the depth of the aragonite saturation horizon, Ωarag = 1.0 (Guinotte et al., 2006). Aragonite saturation state diminishes with depth, due partly to hydrostatic pressure and lower temperature, with a distinct aragonite “saturation horizon” below which waters become under-saturated for aragonite (Ωarag <1) (Jiang et al., 2015). The saturation horizon is a complex outcome of ocean circulation, temperature, CO2 concentrations, salinity, metabolic activity, and the concentrations of organic compounds and occurs at depths between 200 and 3,500 m, depending on the latitude and the ocean (Orr et al., 2005; Doney et al., 2009; Rhein et al., 2013; Jiang et al., 2015). Surface waters and waters at 50 m depth are mostly supersaturated throughout the global ocean (Jiang et al., 2015), however in western Arctic waters, the area of under-saturated waters in the upper 250 m north of 70°N has increased from 5 to 31% between 1990s and 2010 (Qi et al., 2017). At 500 m, large areas of undersaturated Ωarag water are found in the northern and equatorial Pacific ocean. At 1,000 m, Ωarag < 1.8 over all ocean basins and at 2.000 m, Ωarag < 1.0 across all the Pacific and Indian Ocean and parts of the Atlantic Ocean. Ocean acidification is proceeding at higher rates at high latitudes than at lower latitudes (Figure 3) resulting in a shoaling of the aragonite saturation horizon. There is now evidence to show that the aragonite saturation horizon has shoaled since the Preindustrial Period (Turley et al., 2007). For example, in the north east Pacific (from 33.5 to 50.0°N) the aragonite saturation horizon has shoaled by 19.6 m in 11 years (2001–2012) and, at this rate, the entire water column in the northern section of this region is projected to become undersaturated within 50–90 years (Chu et al., 2016).

Biological Responses to a Rapidly Warming and Acidifying Ocean

Not surprisingly, the scale and pace of the physical and chemical changes occurring in the ocean are driving a large range of fundamental responses in marine organisms, ecosystems, and regions (Hoegh-Guldberg et al., 2014; Pörtner et al., 2014). Equally significant, is the observation that relatively small amounts of change have resulted in quite substantial biological impacts, with clear evidence of non-linear trends, tipping points, and otherwise complex responses. Coral responses to changes in ocean conditions, in particular mass coral bleaching, provide particularly compelling examples of the consequences of a rapidly changing ocean for organisms, ecosystems, and dependent societies.

The symbiosis between warm-water corals and Symbiodinium (Figures 4A,B) is very sensitive to changes in the physical and chemical environment surrounding corals. Short periods of high or low temperature and/or light, or exposure to toxins like cyanide, can drive the breakdown of the symbiosis, resulting in the loss of the brown symbionts and a subsequent paling (hence “bleaching”) of the coral host (Hoegh-Guldberg, 1999). Coral bleaching involves the breakdown of the symbiosis between Scleractinian corals and Symbiodinium, which may recover if conditions are not too anomalous for too long. While bleaching of coral tissues has been reported on the scale of colonies or groups of colonies for at least 100 years (Yonge and Nichols, 1931), reports of bleaching at large geographic scales (Figures 4C,D, example of affected coral reefs in American Samoa from late 2015) was unknown to the scientific literature until 1979. Since the early 1980s, however, mass coral bleaching has affected entire reefs and regions, often resulting in significant mortality of reef-building corals. The absence of pre 1979 scientific reports in addition to the close relationship between bleaching and elevated sea temperature, plus considerable laboratory, and mesocosm studies, strongly support the conclusion that mass coral bleaching and mortality are novel and are caused by warm water coral reefs being exposed to rising sea temperatures (Hoegh-Guldberg and Smith, 1989; Glynn, 1993, 2012; Hoegh-Guldberg, 1999; Glynn et al., 2001; Hoegh-Guldberg et al., 2007, 2014; Baker et al., 2008; Eakin C. M. et al., 2010; Strong et al., 2011; Gattuso et al., 2014b). The latest cycle of mass coral bleaching in 2016 (Hoegh-Guldberg and Ridgway, 2016) is reputedly the worst on record and accompanies the warmest years on record (King and Hawkins, 2016;

Figure 4. (A) Scleractinian coral (Turbinaria sp) and (B) Hydrozoan coral (Millepora sp) showing respective Symbiodinium symbionts (each brown cell is about 10 μm in diameter) removed from coral tissues; Credit for (A,B): Todd LaJeunesse, from Pennsylvania State University. ( (C) The photo at left, taken in December 2014, shows coral reef near runway in American Samoa, without obvious bleaching of corals. (D) The photo at right shows the same coral reef, now heavily bleached, in February 2014 (Credit for C,D: Richard Vevers, The Ocean Agency).

Mass coral bleaching and mortality can be triggered by small (1–2°C) SST increases above the long-term summer maxima for a region (Strong et al., 2011). If temperatures are higher for longer, the amount of coral bleaching will increase, driving increased mortality (Hoegh-Guldberg, 1999; Hoegh-Guldberg et al., 2007; Eakin C. M. et al., 2010). There is a strong link between the size and length of temperature extremes and mass coral bleaching and mortality (Hoegh-Guldberg, 1999; Strong et al., 2004, 2011; Eakin C. M. et al., 2010). These relationships are used with satellite data to derive anomalies in SST to monitor the frequency and intensity of mass coral bleaching and mortality (Strong et al., 2004, 2011). For this reason, there is a high level of confidence that the increases in mass coral bleaching and mortality since the early 1980s are due to anthropogenic climate change in particular ocean warming (Hoegh-Guldberg et al., 2014). The loss of symbionts from coral tissues can have immediate effects through the loss of photosynthetic energy, and lead to starvation, disease, reproductive failure, and a loss of competitive ability relative to other organisms on coral reefs (Hoegh-Guldberg and Smith, 1989; Glynn, 1993, 2012; Hoegh-Guldberg, 1999; Baker et al., 2008; Hoegh-Guldberg et al., 2014; Glynn and Manzello, 2015).

Understanding how the positions of ocean isotherms (lines of similar temperatures) are changing and how fast across the ocean surface (“velocity of climate change”, Burrows et al., 2011, 2014) provides insight into whether or not coral populations will be able to move, adapt or acclimatize fast enough to changing sea temperatures (Hoegh-Guldberg, 2012; Pörtner et al., 2014). Some of the highest rates of climate velocity (up to 200 km per decade) were observed in ocean tropical regions (over 1960–2010), driven by shallow spatial gradients in temperature (Burrows et al., 2011, 2014). Observed rates of distribution shifts for individual warm-water coral species linked to increases in sea surface temperatures range from 0 to 150 km per decade, with an average shift rate of 30 km per decade (Yamano et al., 2011; Poloczanska et al., 2013), suggesting that corals and coral ecosystems may be unable to keep up with warming rates (Hoegh-Guldberg, 2012; Burrows et al., 2014; García Molinos et al., 2015).

The possible reduced influence of extremes from climate change with depth has led to the speculation that deeper (>40 m) mesophotic coral reefs may offer a potential refuge against the otherwise rapid changes in temperature, storm intensity, and chemistry that are typical of shallow-water (0–30 m) coral reef environments (Bongaerts et al., 2010a). The “Deep Reef Refugia” hypothesis has been explored by a number of groups who are finding substantial differences in terms of the rate of warming and acidification with depth, as well as examples of species that may span the mesophotic zone to shallow reef areas. Recent work however, has revealed that mesophotic reefs may not be immune to the impacts of storms (Bongaerts et al., 2013). Also, populations of what appear to be the same coral species appear to have considerable genetic structure as a function of depth. This is important given that it implies a high degree of specialization, local adaptation, and even speciation, by corals living at different depths, with the implication that mesophotic corals may not be able to survive in shallow-water environments and vice versa, reducing the potential for mesophotic environments to provide refugia for shallow water Scleractinian corals. This reduces the significance of deeper water populations as a source of recruits for regenerating damaged areas on shallow water coral reefs (Bongaerts et al., 2010b, 2015). In addition to warming oceans, corals are also sensitive to changes to the pH and the carbonate chemistry of seawater as a result of ocean acidification (Kleypas et al., 1999a; Gattuso et al., 2014a). These changes affect organisms in a variety of ways, including reducing calcification rates in a wide array of corals and other organisms in laboratory, mesocosm, and field studies (Gattuso et al., 1998; Reynaud et al., 2003; Kleypas et al., 2006; Dove et al., 2013; Kroeker et al., 2013; Gattuso et al., 2014a).

Long-lived corals from the field have provided an opportunity for retrospective analysis of how growth has varied over long periods of time (De'ath et al., 2009; Lough, 2010, 2011). Calcification measurements from coral cores from 328 colonies of the massive coral Porites growing on the Great Barrier Reef in Australia, for example, have revealed that calcification by these corals has declined by 14.2% since 1990. This appears to unprecedented on the Great Barrier Reef for at least the last 400 years (De'ath et al., 2009) (but see D'Olivo et al., 2013; Figure 5). Given the complexity of the environmental changes occurring in places like the Great Barrier Reef, it is difficult to assign specific drivers of this decline. However, the combined effects of elevated warming and acidification from climate change, along with declining water quality, appear to be significant drivers of the changes observed (D'Olivo et al., 2013). Declining growth and calcification rates have also been detected for Porites colonies in the Red Sea (Cantin et al., 2010) and at several locations in Southeast Asia (Tanzil et al., 2009).

Figure 5. (A–D) Partial-effects plots showing the variation of calcification (grams per square centimeter per year), linear extension (centimeters per year), and density (grams per cubic centimeter) in Porites from the Great Barrier Reef (GBR), Australia, over time. From De'ath et al. (2009). Plots (A–C) are based on 1900–2005 data from 328 Porites colonies, and plot (D) on data for ten long cores. Light blue bands indicate 95% confidence intervals for comparison between years, and gray bands indicate 95% confidence intervals for the predicted value for any given year. Calcification declines by 14.2% from 1990 to 2005 (A), primarily due to declining extension (B). Density declines from 1900 onward (C). The 1572–2001 data show that calcification increased weakly from ~1.62 before 1,700 to ~1.76 in ~1850, after which it remained relatively constant (D) before a weak decline since ~1960. (D–F) Decline coral cover of the GBR over 1985–2012. (E) Map of GBR with color shading indicating mean coral cover averaged over 1985–2014. Points show the location of 214 survey reefs in the northern, central, and southern regions, and their color indicates the direction of change in cover over time. (F) Box plots indicate the percentiles (25, 50, and 70%) of the coral cover distributions within each year and suggest a substantial decline in coral cover over the 27 years. Adapted from De'ath et al. (2012) and with the permission according to PNAS policy.

Studies of the influence of rapidly warming and acidifying conditions on mesophotic coral reefs are absent. Given that these reef systems cover roughly the equivalent area of shallow water coral reefs, understanding how environmental changes are likely to influence these important areas in terms of habitat the fisheries and biodiversity is important and should be a priority of future research (Bongaerts et al., 2010a). Linking the physiological and ecological response of mesophotic reefs to changes in pH and carbonate ion concentration will also be important in the context of understanding how mesophotic coral reef ecosystems will be affected by the shoaling of the saturation horizon in regions such as off Hawaii.

Our understanding of how deep ocean environments are likely to respond to changes in ocean temperature and chemistry are at an early stage. Like mesophotic coral reefs, little is known about the sensitivity of cold-water coral reefs to changes in temperature. As cold-water corals tend not to have a mutualistic symbiosis with Symbiodinium, their response is naturally different to that of symbiotic Scleractinian corals. As with mesophotic coral reefs, there is much more to be discovered with respect to how these critically important cold-water coral reefs are likely to respond to steadily warming and acidifying ocean. Coral reefs in the deep-sea have been identified as particularly vulnerable to the effects of ocean acidification; in part because of the numerical predominance of calcifying taxa, and in part because the pre-industrial carbonate levels at the depths and temperatures they inhabit were already low (Freiwald et al., 2004). Experimental studies reveal that short-term exposures of important deep-water corals such as L. pertusa to a reduction in pH of around 0.15–0.3 units resulted in a decrease in calcification rates of between 30 and 56% (Maier et al., 2009). However, subsequent work has shown that L. pertusa can acclimatize (i.e., maintain considerable calcification) to declining aragonite levels modifying skeletal structure and skeletal strength (Form and Riebesell, 2012; Hennige et al., 2015). Observations of deep-sea corals in under-saturated waters from the SW Pacific also suggest some species-specific tolerance, however growth rates are extremely low and in under-saturated conditions dead coral skeletons dissolve rapidly (Bostock et al., 2015; Thresher et al., 2015). Whether cold-water corals will be able to adjust to rapid warming and ocean acidification projected for the coming century is unknown. However, analyses of cold-water coral fossils suggest that a combination of declining aragonite and oxygen saturations will reduce the distribution of cold-water corals (Thiagarajan et al., 2013).

While coral species and their symbionts have received a major amount of focus in terms of the effect of ocean warming and acidification on warm-water coral reef ecosystems, there is a growing number of studies that have revealed impacts on a broader range of reef organisms. Among the most affected are calcifying algae, calcareous phytoplankton, molluscs, and echinoderms, with the larval stages of some organisms being more sensitive than the adult phase (Kroeker et al., 2013). Bioeroding organisms also respond to both warmer and more acidic conditions (Dove et al., 2013; Fang et al., 2013; Reyes-Nivia et al., 2013). The sponge, Cliona orientalis, increased biomass and bioerosion capability when exposed to warmer and more acidic conditions, implicating a role of this sponge in helping tip the carbonate balance of reefs toward net erosion (Dove et al., 2013; Fang et al., 2013). Similar observations have been made for bio-eroding endolithic algal communities, where small shifts in ocean temperature and acidity (i.e., CO2 levels) enhanced skeletal dissolution and was associated with increased endolithic biomass and respiration under elevated temperatures and CO2 levels (Reyes-Nivia et al., 2013).

OA Observations and Data

Follow the links below to access ocean acidification data for each of our observation programs

The field of carbon cycle science depends on well-designed, well-executed, and carefully maintained observations.  The PMEL carbon group primarily focuses on large scale observations of ocean interior carbon through hydrographic cruises and surface ocean carbon dynamics through measurements made on volunteer observing ships, buoys, and other autonomous systems. We work in both the open ocean and in coastal environments. We maintain long-term time series observations as well as conducting short term process studies or exploratory studies.  Since ocean acidification emerged as an important scientific issue, we have been augmenting and expanding our observational capacity by adding pH and other biogeochemical measurements to the platforms listed below.

Volunteer Observing Ships (VOS)

For the last 2 decades, we have used underway sampling on research vessels and VOS to measure large-scale trends in ocean carbon chemistry.  We are in the process of adding pH and additional parameters necessary to address ocean acidification using VOS. 

Buoys and Autonomous Systems

High frequency observations provide the basis to better understand natural variability of ocean acidification over daily to seasonal cycles. We are currently measuring pCO2 and pH on 19 moorings in open ocean, coastal U.S., and coral reef waters. Click the links below for more information.


Open Ocean

Coral Reef

Hydrographic Cruises

Ship-based hydrography is the only method for obtaining high-quality, high spatial and vertical resolution measurements of a suite of physical, chemical, and biological parameters over the full water column, and in areas of the ocean inaccessible to other platforms.  Our hydrographic cruises in the North Pacific Ocean, along the U.S. West Coast (2007, 2011, 2012, 2013) and in the Puget Sound have revealed a trend of ocean acidification over time, led to a better understanding of the spatial and temporal variability of ocean acidification, and provided insights into how ocean acidification will manifest in estuaries and near shore coastal areas.


See the links below to learn more about ocean acidification and the type of research our group is involved in.

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