BGBiogeosciencesBGBiogeosciences1726-4189Copernicus PublicationsGöttingen, Germany10.5194/bg-13-1767-2016Reviews and Syntheses: Ocean acidification and its potential impacts
on marine ecosystemsMostofaKhan M. G.mostofa@tju.edu.cnhttps://orcid.org/0000-0001-5203-2432LiuCong-Qiangliucongqiang@vip.skleg.cnZhaiWeiDonghttps://orcid.org/0000-0001-9410-1045MinellaMarcoVioneDavideGaoKunshanhttps://orcid.org/0000-0001-7365-6332MinakataDaisukeArakakiTakemitsuYoshiokaTakahitoHayakawaKazuhideKonohiraEiichiTanoueEiichiroAkhandAnirbanChandaAbhraWangBaoliSakugawaHiroshiInstitute of Surface-Earth System Science, Tianjin University, Tianjin 300072, ChinaState Key Laboratory of Environmental Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang 550002, ChinaInstitute of Marine Science and Technology, Shandong University, Ji-nan 250100, ChinaUniversità degli Studi di Torino, Dipartimento di Chimica, Via P. Giuria 5, 10125 Torino, ItalyCentro Interdipartimentale NatRisk, Via Leonardo da Vinci 44, 10095 Grugliasco (TO), ItalyState Key Laboratory of Marine Environmental Science (B-606), Xiamen University, Daxue Rd 182, Xiamen, Fujian 361005, ChinaDepartment of Civil and Environmental Engineering, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931, USADepartment of Chemistry, Biology and Marine Science, Faculty of Science, University of the Ryukyus, Senbaru, Nishihara-cho, Okinawa 903-0213, JapanInstitute for Hydrospheric–Atmospheric Sciences, Nagoya University, Nagoya, JapanLake Biwa Environmental Research Institute, Shiga Prefecture, Ohtsu 520-0806, JapanHydrospheric Atmospheric Research Center, Nogoya University, Nagoya, JapanSchool of Oceanographic Studies, Jadavpur University, Jadavpur, Kolkata 700032, West Bengal, IndiaGraduate School of Biosphere Science, Department of Environmental Dynamics and Management,
Hiroshima University, 1-7-1 Kagamiyama, Higashi-Hiroshima 739-8521, JapanPresent address: Field Science Education and Research Center, Kyoto University, Kitashirakawa Oiwake-cho, Sakyo-ku, Kyoto 606-8502, JapanPresent address: DLD inc., 2435 Kamiyamada, Takatomachi, Ina, Nagagano 396-0217, JapanKhan M. G. Mostofa (mostofa@tju.edu.cn) and Cong-Qiang Liu (liucongqiang@vip.skleg.cn)23March20161361767178622June201513July201515January201612March2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://bg.copernicus.org/articles/13/1767/2016/bg-13-1767-2016.htmlThe full text article is available as a PDF file from https://bg.copernicus.org/articles/13/1767/2016/bg-13-1767-2016.pdf
Ocean acidification, a complex phenomenon that lowers seawater pH,
is the net outcome of several contributions. They include the dissolution of
increasing atmospheric CO2 that adds up with dissolved inorganic
carbon (dissolved CO2, H2CO3, HCO3-, and
CO32-) generated upon mineralization of primary producers (PP) and
dissolved organic matter (DOM). The aquatic processes leading to inorganic
carbon are substantially affected by increased DOM and nutrients via
terrestrial runoff, acidic rainfall, increased PP and algal blooms,
nitrification, denitrification, sulfate reduction, global warming (GW), and
by atmospheric CO2 itself through enhanced photosynthesis. They are
consecutively associated with enhanced ocean acidification, hypoxia in
acidified deeper seawater, pathogens, algal toxins, oxidative stress by
reactive oxygen species, and thermal stress caused by longer stratification
periods as an effect of GW. We discuss the mechanistic insights into the
aforementioned processes and pH changes, with particular focus on processes
taking place with different timescales (including the diurnal one) in
surface and subsurface seawater. This review also discusses these collective
influences to assess their potential detrimental effects to marine organisms,
and of ecosystem processes and services. Our review of the effects operating
in synergy with ocean acidification will provide a broad insight into the
potential impact of acidification itself on biological processes. The
foreseen danger to marine organisms by acidification is in fact expected to
be amplified by several concurrent and interacting phenomena.
Introduction
Ocean acidification is typically defined as a process of
increasing seawater acidity or lowering seawater pH, as a consequence of the
dissolution of elevated atmospheric CO2. Carbon dioxide from the
atmosphere (Orr et al., 2005; Feely et al., 2008) adds to the dissolved
inorganic carbon (DIC; dissolved CO2, H2CO3,
HCO3-, and CO32-) originated from the degradation of
dissolved organic matter (DOM) (Mostofa et al., 2013a), primary producers
(PP) (Cai et al., 2011; Mostofa et al., 2013a), CO2 seeps from
sub-seabed storage (Taylor et al., 2014) and volcanic vents (Lidbury et
al., 2012; Hall-Spencer et al., 2008) in shallow submarine zones, anaerobic
oxidation of methane (Haroon et al., 2013), and sulfide oxidation coupled to
carbonate dissolution (Torres et al., 2014) in seawater. The sources of
elevated atmospheric CO2 first of all include anthropogenic
activities such as fossil fuels combustion, i.e., coal, petroleum, and
natural gas (Le Quéré et al., 2009), enhanced land-use practices (Le
Quéré et al., 2009), as well as deforestation (van der Werf et
al., 2009; Lapola et al., 2014). Additionally, there could be significant
contributions from natural sources such as plant litter decomposition (King
et al., 2012), volcanic eruptions (Hall-Spencer et al., 2008), emission of
CO2 from freshwater including the Amazon River basin (Sobek et
al., 2005; Abril et al., 2014), and enhanced respiration of soil organic
matter (OM) under global warming (GW) conditions (Knorr et al., 2005).
The emissions of CO2 by fossil fuels combustion increased by
29 % in 2000–2008 (Le Quéré et al., 2009) and, as far as
natural-water sources are concerned, the contribution from European estuaries
is for instance equivalent to approximately 5–10 % of the anthropogenic
CO2 emissions in western Europe (Frankignoulle et al., 1998).
Recent studies demonstrate that ocean acidification under elevated
CO2 and temperature levels could increase primary productivity of
specific species (Holding et al., 2015; Coello-Camba et al., 2014; Li et
al., 2012). Additionally, such specific species-based primary productivity is
also found to increase either by an increasing seawater CO2 level (Kim
et al., 2006; Olischläger et al., 2013) or elevated temperature alone
because of the effects of global warming (Yvon-Durocher et al., 2015;
Lewandowska et al., 2012). The primary production in the oceans contributes
approximately 48.5 petagrams (1 Pg=1015g) of
Cyr-1 (46.2 % of the total), as estimated using the
integrated the Carnegie–Ames–Stanford Approach – the Vertically Generalized Production Model (CASA-VGPM) biosphere model (Field et al., 1998). As a consequence,
approximately one-third to 50 % of the atmospheric CO2 is fixed
annually worldwide by marine phytoplankton (Sabine et al., 2004; Toseland et
al., 2013). However, one should also consider that the photoinduced and
biological mineralization of organic matter (OM), including DOM and dead
organisms, is an important source of DIC in seawater and liberates again an
important fraction of the CO2 fixed by photosynthesis (Bates and
Mathis, 2009; Mostofa et al., 2013a).
Ocean acidification is responsible for changes in the oceanic carbonate
system, with effects on partial pressure of CO2
(pCO2), DIC, pH,
alkalinity, and calcium carbonate saturation state (Feely et al., 2010;
Beaufort et al., 2011). In the case of calcifying organisms one observes a
marked pattern of decreasing calcification with increasing pCO2,
which follows the corresponding decreasing concentrations of CO32-
as a consequence of decreasing pH (Beaufort et al., 2011). Such effects
finally cause a decline in calcification and growth rates of shellfish
(Talmage and Gobler, 2010; Wittmann and Pörtner, 2013), of shell-forming
marine plankton and of benthic organisms including corals (Kleypas et
al., 1999; Doney et al., 2009; Beaufort et al., 2011; Pandolfi et al., 2011;
McCulloch et al., 2012). The latter have already been lost or are highly
damaged in coastal areas near many countries including Indonesia, Hawaii,
the Caribbean, Fiji, Maldives, and Australia (Erez et al., 2011). A 30 %
decline or damage of coral reef ecosystems has been estimated worldwide, and
it is predicted that as much as 60 % of the world's coral reefs might be
lost by 2030 (Hughes et al., 2003).
A conceptual model of acidification in coastal to open oceans,
showing either dissolution of atmospheric CO2 or emission of aquatic
CO2 plus DIC originated from the photoinduced and/or biological
respiration of primary producers (PP). The latter includes both dissolved
organic matter (DOM) and PP (1). Uptake of such CO2 is primarily
responsible for the occurrence of photosynthesis and PP (2) that can generate
algal toxins or pathogens in the euphotic zone, along with generation of
CO2, DIC, and other products; PP can also be enhanced by autochthonous
DOM (2), by DOM or sinking cells in subsurface or deeper seawater (2), and by
riverine DOM (2). Atmospheric acid rain (mostly HNO3 and
H2SO4) can contribute directly to the acidification (3). Global
warming can lengthen the stratification period with a subsequent decline in
vertical mixing, which reduces the exchange with surface oxygenated
water (4).
The extent and effects of ocean acidification can be exacerbated by several
complex processes, some of which act as stimulating factors, such as local
environmental impacts including terrestrial or riverine runoff (Sunda and
Cai, 2012; Bauer et al., 2013), modified land use practices (Lapola et al.,
2014), and atmospheric acid rain (Baker et al., 2007). An additional effect
could be represented by the enhanced mineralization of DOM and PP (e.g.,
phytoplankton) as a consequence of global warming (Mostofa et al., 2013a).
Such mineralization could be biological (respiration) or abiotic via
different (mainly) photochemical processes. Most of the cited effects are
expected to cause eutrophication or algal blooms in coastal seawater, which
would in turn affect the carbon cycling and the carbonate chemistry and
influence the overall acidification process (Beaufort et al., 2011; Sunda and
Cai, 2012; Bauer et al., 2013). Such acidification is responsible for changes
in the oceanic carbonate system (Feely et al., 2010; Beaufort et al., 2011),
which subsequently impacts on marine living organisms and the related
ecosystem processes or services (Cooley et al., 2009; Mora et al., 2013;
Mostofa et al., 2013a). Considering the possible devastating consequences on
the marine ecosystems, their organisms and the related ecosystem services
(Cooley et al., 2009; Doney et al., 2009, 2012; Cai, 2011), it is important
to ascertain all the possible causes of ocean acidification and their
interlinks.
This review will provide a general overview of the ocean acidification such
as chemistry and ecological consequences, including the interactions between
acidification by CO2 and other processes that could in turn modify the
seawater pH. We shall discuss changes in the pH values in both sea surface
and subsurface/deeper water extensively with different timescales, from
diurnal to multi-annual. We shall also address potential impacts of ocean
acidification on marine organisms, along with possible indirect impact
processes from a series of stimulating factors (oxidative stress in surface
seawater, hypoxia in subsurface/deeper seawater, stress caused by algal, or
red-tide toxins and pathogens) for both sea surface and subsurface/deeper
water. Our review from the point of synergistic effects of ocean acidification
with such stimulating factors will broaden the understanding of the potential
impact of acidification on biological processes. Such an impact is based on the
conceptual model provided for both surface and deeper seawaters.
Potential mechanisms behind ocean acidification
Ocean acidification includes several potential phenomena that may be
operational at the global and/or local scales (Fig. 1): (i) increasing
dissolution of atmospheric CO2 to seawater: anthropogenic ocean
acidification; (ii) input of CO2 plus DIC upon mineralization of PP
influenced by elevated atmospheric CO2: natural ocean acidification;
(iii) enhanced PP and respiration due to the effects of global warming and
other processes: natural ocean acidification; and (iv) direct acidification
and stimulation of PP by atmospheric acid rain: natural and anthropogenic
ocean acidification. A pictorial scheme of the main operational processes
affecting the ocean acidification is depicted in Fig. 1.
Increasing dissolution of atmospheric CO2 to seawater:
anthropogenic ocean acidification
Enhanced dissolution of atmospheric CO2 to seawater lowers pH and
modifies the carbonate chemistry, affecting both biogenic and sedimentary
CaCO3. This process has extensively been discussed in earlier reviews
(Pearson and Palmer, 2000; Feely et al., 2008; Beaufort et al., 2011). For
the given seawater, net CO2 fluxes (either from atmosphere to water
or the reverse) may significantly vary depending mostly on time (day or
night) and season. Based on a series of studies, six scenarios can be
formulated for the net sea–air fluxes of CO2. They are (i) sinking
or balance of atmospheric CO2 to seawater under sunlight, and
emission or balance of CO2 to the atmosphere during the night;
(ii) emission or balance of CO2 to the atmosphere during daytime, and
sinking or balance of atmospheric CO2 to surface water during the
night; (iii) emission or balance of seawater CO2 to the atmosphere
during both day and night; (iv) sinking or balance of atmospheric CO2
to surface water during both day and night; (v) sinking or source or balance
of atmospheric CO2 to surface water during the warm period; and
(vi) emission or sinking or balance of seawater CO2 to the atmosphere
during the cold period. These scenarios are described in the Supplement.
Diurnal variation of pH along with fCO2[seawater]
(µatm) and fCO2[air] (µatm) in surface
seawater of the Jiulongjiang estuary (a) and the Bay of
Bengal (b). pH, fCO2[seawater] (µatm), and sea
subsurface temperature for seawater samples (from 13 to 75 m depth) in the
northern Yellow Sea (c). Samples from the Jiulongjiang estuary were
collected from 28 June 2009 at 16:00 LT (Chinese Standard Time) to 29 June
2009 at 14:55 LT, from 24∘25′ to 24∘46′ N and
118∘00′ to 119∘19′ E. Throughout the sampling
period (a) there was a range of salinity (4.4–33.9 psu) and
of sea surface water temperature (26.59–29.12 ∘C). Samples from the
Bay of Bengal were collected in May whereas pH, fCO2[seawater],
and fCO2[air] varied from 8.12 to 8.37, 153 to 373, and 370 to
381 µatm, respectively, along with salinity (27.82±0.26psu), chlorophyll a (12.35±2.23µgL-1),
sea surface water temperature (SST: 28.50–31.70 ∘C), and daytime
solar intensity flux (556–109700–17 Lux at 05:00, 12:00, and 18:00 LT,
respectively) (b). The calculated pH, salinity,
fCO2[seawater] at in situ DIC, and SST varied in the respective
ranges 7.53–7.97, 28.24–32 psu, 280–776 µatm, and
3.44–20.58 ∘C, respectively, for the subsurface samples collected from northern
Yellow Sea with the range
37∘25′–39∘67′ N latitude and
121∘16′–124∘10′ E longitude (c).
Input of CO2 plus DIC upon mineralization of PP influenced by elevated atmospheric
CO2: natural ocean acidification
The formation and seawater dissolution of CO2 and DIC produced from
photoinduced and biological mineralization of PP or
DOM also lowers pH and modifies the carbonate
chemistry (Fig. 2) (Cai et al., 2006, 2011; Feely et al., 2010; Sunda and
Cai, 2012; Bates et al., 2013; Mostofa et al., 2013a). Anticorrelation
between pH and CO2 levels during the diurnal cycle has been observed
in surface and sub-surface waters (Fig. 2), where CO2 is mainly
originated from the biological respiration of PP or DOM. Such an issue is
further complicated by the fact that enhanced levels of CO2 are
partially responsible for the increase of photosynthesis (Behrenfeld et
al., 2006; Kranz et al., 2009), and they may have a deep impact on the net
primary production (PP) (Hein and Sand-Jensen, 1997; Behrenfeld et al., 2006;
Jiao et al., 2010). The upper ocean organisms, mostly the autotrophs, are a
massive carbon-processing machine that can uptake atmospheric CO2
(Hein and Sand-Jensen, 1997; Falkowski et al., 1998; Sarmento et al., 2010)
or CO2 plus DIC regenerated from DOM or PP, particularly during the
daytime (Fig. 2a; see also Supplement) (Takahashi et al., 2002, 2009; Yates et al., 2007; Chen and
Borges, 2009; Mostofa et al., 2013a). In contrast, during the night seawater
can become a source of CO2, as shown in Fig. 2 in three different
contexts. The ability of water to act as a CO2 source is shown by the
higher values of pCO2 in seawater compared to that in the atmosphere
(Zhai et al., 2005, 2014; Yates et al., 2007; Chen and Borges, 2009).
The daytime uptake of CO2 is the consequence of primary production
through photosynthesis, which mostly uses dissolved CO2 via the
enzyme ribulosebiphosphate carboxylase (RUBISCO), which governs the
carbon-concentrating mechanisms (CCMs) (Yoshioka, 1997; Behrenfeld et al.,
2006; Kranz et al., 2009). Mesocosm experiments using 14C-bottle
incubations indicate that elevated CO2 can increase 14C primary
production or bacterial biomass production, also leading to the formation of
dissolved organic carbon (DOC) and to its rapid utilization (Engel et
al., 2013).
Photosynthetic carbon fixation by marine phytoplankton leads to the formation
of ∼ 45 gigatons of organic carbon per annum, of which 16 gigatons
(∼ 35.6 % of the total) are exported to the ocean depths (Falkowski
et al., 1998). Furthermore, all primary producers including the large and
small cells can contribute to the carbon export from the surface layer of the
ocean, at rates proportional to their production rates (Richardson and
Jackson, 2007). The reprocessing of this organic material can cause a
decrease in the pH of seawater via the CO2 produced by respiration
(Jiao et al., 2010). If, in addition, organic N and P are biologically
transformed into NO3- and phosphate (Mostofa et al., 2013a) and if
there is also transformation of NH4+ to N2 (Doney et
al., 2007), there can be a further decrease of seawater alkalinity. Such
processes also decrease the buffering capacity of seawater (Thomas et
al., 2009), which would become more susceptible to acidification caused by
the dissolution of atmospheric CO2 (Thomas et al., 2009; Cai et
al., 2011). A decrease in alkalinity and accompanying acidification may have
negative impacts on shellfish production (Hu et al., 2015).
Heterotrophic bacteria are the main organisms that are responsible for
respiration in the ocean (> 95 %) (Del Giorgio and Duarte, 2002), and
half of the respiration (approximately 37 Gt of C per year) takes
place in the euphotic layer (del Giorgio and Williams, 2005). An interesting
issue is that such bacteria are also important sources of the superoxide
radical anion (O2-•) (Diaz et al., 2013), the
dismutation of which (2O2-•+2H+→H2O2+O2) consumes H+ and could partially buffer
at local scale the acidification that is connected to the degradation of OM
(Mostofa et al., 2013b).
The biological transformation of DOM and PP is active constantly at the sea
surface as well as in the subsurface/deeper water, whilst photoinduced
degradation is merely active during daytime in the sea surface layer. Of
course, such processes show variations associated with seasonal and annual
changes in deep-sea geochemistry and biology, along with phenomena associated
with ocean circulation (Asper et al., 1992; Thomas et al., 2004). The entire
phytoplankton biomass of the global oceans is consumed every 2 to 6 days
(Behrenfeld and Falkowski, 1997) and part of the carbon fixed by the
autotrophs is actually respired in situ (Sarmento et al., 2010), also
providing nutrients for the microbial food web (Behrenfeld et al., 2006;
Sarmento et al., 2010). In some cases, the reprocessing of nutrients is
involved in harmful algal blooms or eutrophication by enhanced photosynthesis
in surface seawater (Sunda and Cai, 2012; Mostofa et al., 2013a).
Enhanced PP and respiration due to the effects of global warming
and other processes: natural ocean acidification
Anthropogenic global warming could also enhance the natural acidification
process. The dissolution of CO2(g) and DIC released from PP and its
subsequent respiration/degradation can be enhanced by the effects of GW
(Behrenfeld et al., 2006; Cai et al., 2006; Kranz et al., 2009; Cai et al.,
2011; Sunda and Cai, 2012; Mostofa et al., 2013a; Holding et al., 2015). GW
is a key factor to increase water temperature (WT), which can affect the
extent and the duration of the vertical stratification during the summer
season. Furthermore, the prolonged exposure of the surface water layer to
sunlight may cause photoinduced bleaching of sunlight-absorbing DOM, the
so-called color dissolved organic matter (CDOM), thereby enhancing the water
column transparency and modifying the depth of the mixing layer or euphotic
zone (Behrenfeld et al., 2006; Huisman et al., 2006). The increased stability
of the water column may also enhance the photoinduced and biological
mineralization of OM, due to the combination of higher temperature and of the
longer exposure of the water surface layer to sunlight (Huisman et al., 2006;
Vázquez-Domínguez et al., 2007). A further effect is the reduction
of subsurface dissolved O2 because of the decline of vertical winter
mixing, which subsequently reduces the exchange of surface oxygenated water
to the deeper layers (Fig. 1). Increasing temperature increases the
respiration rates in natural waters (Vázquez-Domínguez et
al., 2007), and it affects phytoplankton metabolism nearly as significantly
as nutrients and light do (Toseland et al., 2013). Various photoinduced and
microbial products/compounds formed from DOM or PP (e.g., CO2, DIC,
H2O2, NH4+, NO3-, PO43-, CH4,
autochthonous DOM), the generation of which can be higher in stratified
surface water as a consequence of GW, may enhance photosynthesis and,
consequently, primary production as schematized in Fig. S1 in Supplement
(Bates and Mathis, 2009; Cai et al., 2011; Mostofa et al., 2013a). Further
details are reported in the Supplement.
Diurnal changes of pH, H2O2 and solar (UV) intensity in the
seawater of Taira Bay on 9–10 January 2003 (a) and Sesoko Bay on
19–20 January 2003 (b). (c, d): pH and concentration of
H2O2 as a function of the solar UV intensity with the related linear
fit regressions in the case of Taira Bay and Sesoko Bay samples,
respectively. In the seawater of Taira Bay the pH, H2O2, dissolved
organic carbon (DOC), and sea surface water temperature (SST) varied in the
following ranges: 8.16–8.25, 40-100 nM, 1.14–1.42 ppm, and
18.8–20.9 ∘C, respectively. In the seawater of Sesoko Bay the
relevant ranges were as follows: 7.82–8.28, 30–110 nM,
0.84–1.41 ppm, and 17.7–20.2 ∘C, respectively.
Diurnal, abrupt and homogeneous pH changes in seawater
In some locations, the pH of the sea surface water gradually increases during
the period before sunrise to noon and then decreases after sunset as a
function of the solar irradiation flux (Fig. 3a, b) (Fransson et al., 2004b;
Arakaki et al., 2005; Akhand et al., 2013). Furthermore, substantial
fluctuations of the pH values during daytime are also observed (Fig. 3a, b)
(Fransson et al., 2004a; Arakaki et al., 2005; Clark et al., 2010). The
magnitude of the diurnal pH variation can be substantial, ranging from
∼ 0.01 in waters with low biological activity to 1.60 in waters with
high biological activity that are influenced by riverine inputs, particularly
in coastal areas (Table S1 in Supplement). More specifically, pH has been
observed to increase by 0.03–0.81 units in surface coastal seawater, from
0.26 to 1.60 in macroalgae, 0.01 to 0.75 in coral reefs, from 0.17 to
> 1.00 in the seagrass community, from 0.03 to 1.59 in CO2 venting
sites, and from 0.04 to 0.10 in polar oceans (Table S1) (Semesi et al., 2009;
Taguchi and Fujiwara, 2010; Hofmann et al., 2011). Diurnal pH changes in sea
surface waters are apparently triggered by two phenomena. The first and key
issue is the consumption or dissolution in seawater of CO2 that is
involved in primary production (Fig. 2a, b) (Akhand et al., 2013; Zhai et
al., 2014). Depending on the ratio between photosynthesis and respiration,
diurnal fluctuations of pCO2 are observed in seawater and the
pCO2 maxima correspond to pH minima and vice versa. In the case of
Fig. 2a and b, the pH maxima are observed at noon or soon after noon; in
other locations they may occur in different times of the day, but the
anticorrelation between pH and pCO2 is always observed. At the sea
surface one may observe a diurnal decrease in pCO2 with an increase
in pH during the daytime or in the presence of sunlight (due to the
prevalence of photosynthesis), along with an increase in pCO2 with
a decrease in pH at night when respiration prevails (Yates et al., 2007; Semesi
et al., 2009).
A second issue that might affect pH is the photoinduced generation of
H2O2, primarily by dismutation of superoxide radical anion
(2O2•-+2H+→H2O2+O2)
(Fig. 3a, b) (Arakaki et al., 2005; Clark et al., 2010) and the subsequent
production of the strong oxidant, hydroxyl radicals (HO•)
via photolysis or Fenton and photo-Fenton processes, which are responsible
for the degradation of DOM and POM (Vione et al., 2006; Minakata et
al., 2009). The linear correlation between pH / [H2O2] and the
UV (ultraviolet) intensity (Fig. 3c, d) can be elucidated by considering that both
variables are directly influenced by solar irradiation.
Seawater pH is predominantly determined by the balance between consumption
(photosynthesis) and release (respiration) of CO2 as a consequence of
the PP activity. In the reported cases the maximum consumption of dissolved
CO2 takes place at the same time as the maximum activity of the
photo-stimulated biota. In addition, the positive correlation between
[H2O2] and UV intensity (Fig. 3c, d) is linked to the fact that the
O2-• production rate overlaps with the maximum of solar
irradiation, because the biological and photochemical production of
O2-• is activated by light absorption. The concentration
of H2O2 in sea surface water gradually increases during the period
before sunrise to noon and then decreases after sunset as a function of solar
irradiation (Fig. 3a, b). The amplitude of the H2O2 diurnal cycle
(the highest concentration at noontime minus concentration during the period
before sunrise) ranged from 20 to 365 nM in coastal seas to marine
bathing waters (Table S1). Both the O2-• production and
its dismutation with formation of H2O2 involve H+ exchange
and can consequently affect the ocean pH. O2-• is
largely produced by the enzyme nicotinamide adenine dinucleotide phosphate (NADPH) oxidase through the synthesis of
HO2• that is a weak acid (pKa = 4.88) (Bielski et
al., 1985), which dissociates at the oceanic pH releasing H+ ions
according to the following reactions:
O2+NADPH→NADP+HO2•,HO2•O2-•+H+.
The production and dismutation of O2-• is a
H+-neutral process, but the fate of the superoxide anion is also a
consequence of the redox state of the environment. Indeed, superoxide can be
oxidized to O2 (O2-•→O2+e-) or reduced to H2O2 (O2-•+e-+H+→H2O2). The prevalence of one of the two processes
may not have the same effect on the overall H+ budget and can
consequently affect the acid-base equilibria of oceanic seawater. The
generation of O2-• and consequently of H2O2
(Fig. 3a, b) would give an additional contribution to the daytime pH maxima
and, as a consequence, could be a further actor in the definition of the
daytime pH fluctuation.
Apart from the diurnal cycle, abrupt pH changes caused by both photoinduced
and biological processes (overlapping to diurnal changes) have been observed
in surface seawater and among the branches of Pocillopora colonies
in the Great Barrier Reef (Gagliano et al., 2010), in the surface seawater of
Okinawa Island (Fig. 3a, b) (Arakaki et al., 2005), in marine bathing waters
(southern California) (Clark et al., 2010), in the North Sea (Blackford and
Gilbert, 2007), in the North Pacific Ocean (Byrne et al., 2010), in the
Chwaka Bay (Semesi et al., 2009), and in the northeast Atlantic (Findlay et
al., 2014). Such rapid changes in pH are supposed to be a consequence of the
primary production as well, although the details of the pH-modifying
pathway(s) are still poorly understood. Proposals include several processes
in which an intracellular microenvironment is produced, with very different
pH values compared to the surrounding seawater, with possible release of
intracellular material as a consequence of, e.g., cell lysis. Among these
processes the main ones are (i) pH variation connected with aggregates present in
photosynthetically active cells or inside colonies (Lubbers et al., 1990);
(ii) polyanion-mediated formation of mineral–polymer composites inside
alginate microgels or in the Golgi of coccolithophorid algae (Chin et
al., 1998); (iii) processes occurring at the site of calcification such as
conventional H+-channeling, Ca2+–H+ exchanging
ATPase, transcellular symporter, and co-transporter H+-solute
shuttling (Ries, 2011); (iv) cellular extrusion of hydroxyl ions
(OH-) into the calcifying medium (Ries, 2011); and
(v) CO2 consumption via photosynthesis (Ries, 2011). Furthermore, the
ability to up-regulate pH at the site of calcification can provide corals
with enhanced resilience to the effects of ocean acidification (McCulloch et
al., 2012). Increased pH during high primary productivity can be justified by
the observation of a parallel increase in the δ13C values of POM,
which may reflect a shift by phytoplankton from using CO2 to using
HCO3- for photosynthesis (Doi et al., 2006; Akhand et al., 2013).
Therefore, uptake of HCO3- for phytoplankton photosynthesis at
high pH might be the effect of its enhanced occurrence in seawater.
Homogeneous (longer-term and constant-rate) acidification in
subsurface/deeper seawater is characteristically observed in oceans (Fig. 2c,
Table S1; Feely et al., 2008; Byrne et al., 2010; Taguchi and Fujiwara, 2010;
Cai et al., 2011; Zhai et al., 2012; Bates et al., 2013), estuaries (Feely et
al., 2010), and experimentally in dark incubation (Lubbers et al., 1990).
Such a homogeneous pH behavior is also followed in the subsurface water of a
large freshwater lake (Fig. S2a). At the beginning of the summer
stratification period, pH in subsurface water (at depths of 40 and
80 m) gradually decreases whilst pH in the surface lake water (at
depths of 2.5 and 10 m) increases, while DOC (Fig. S2b) and PP (chlorophyll a, Fig. S2c) also increase. Similar
results, particularly monthly pH variations in surface and deeper seawater,
are observed in the Seto Inland Sea during the summer stratification period
and during convective mixing periods (Taguchi and Fujiwara, 2010).
Homogeneous acidification can vary on a timescale of days to weeks or even
months in a wide range of subsurface water at a specific depth (Fig. S3;
Byrne et al., 2010; Taguchi and Fujiwara, 2010). For example, pH was 7.4 at
∼ 2000–2500 m depth and 7.5 at ∼ 2500–3400 m depth along
25–55∘ N in the North Pacific Ocean (Byrne et al., 2010), or pH was 7.0
at 80 m depth during the August–November period (Fig. S3). In the dark, pH
decreases gradually inside colonies and also “nightly” decreases of pH
occur (Lubbers et al., 1990). Such homogeneous acidification is primarily
linked to the dissolution of CO2 plus DIC originated from the
biological degradation of sinking microorganisms (Bates and Mathis, 2009; Cai
et al., 2011) and of the DOM originally produced by such organisms (Mostofa
et al., 2013a). Enhanced acidification due to the biological degradation of
OM can cause undersaturation of aragonite and calcite during the summer
period in subsurface/deeper seawater in the Yellow Sea (Figs. 2b, 4) (Zhai et
al., 2013), Gulf of Mexico (Cai et al., 2011; Sunda and Cai, 2012), North
Pacific Ocean (Byrne et al., 2010), Arctic Ocean (Bates et al., 2013), and
Arctic shelves (Bates and Mathis, 2009).
The biological degradation processes are constantly occurring in
subsurface/deeper seawater after the onset of early summer, and they continue
during the summer stratification period for several months, until the start
of winter vertical mixing (Fig. 1). The occurrence and importance of these
processes is shown by the increasing trend in subsurface CO2 followed
by a similar decreasing trend of pH. Significant anticorrelation between the
two parameters (r2=0.5) has been observed in subsurface seawater
(13–75 m depth) along 37∘25′–39∘67′ N to
121∘16′–124∘10′ E in the Yellow Sea (Fig. 4a).
Furthermore, the same evidence was observed in the Seto Inland Sea (Taguchi
and Fujiwara, 2010) and in the diurnal samples of the Luhuitou fringing reef
(Sanya Bay) of the South China Sea (Zhang et al., 2013). Strong anticorrelation
between pCO2[seawater] and dissolved O2 (r2=0.8;
Fig. 4b) supports the production of CO2 plus DIC from the biological
respiration/degradation of DOM and PP by heterotrophic bacteria as discussed
earlier. Such bacteria also produce the superoxide radical anion
(O2•-) (Diaz et al., 2013) that might be further
involved in the processing/oxidation of DOM or PP by producing H2O2
and consequently •OH via photolysis, photo-Fenton, or Fenton-like
processes. Such trends of CO2 (or DIC) vs. dissolved O2 are
also observed in California coastal waters (DeGrandpre et al., 1998), in the East
China Sea (Zhai and Dai, 2009), in the South China Sea (Zhai et al., 2009), and
in Seto Inland Sea (Taguchi and Fujiwara, 2010). Biological respiration can
be evidenced from an experiment conducted using subsurface water (37 m
depth) collected from the East China Sea, where the decline in dissolved
O2 is significantly coupled with an increase of DIC production during
a 60-hour study period (Fig. 4c). The heterotrophic bacteria carry out the
largest fraction of respiration (> 95 %) in the ocean (Del Giorgio and
Duarte, 2002). This means that the heterotrophic community catabolizes an
important percentage of the OM produced by the autotrophs (e.g., plants,
algae,
or bacteria) (Laws et al., 2000). Therefore, enhanced primary production or
algal blooms in surface seawater and the subsequent sinking are the key
processes for homogeneous acidification of the subsurface layer during the
summer stratification period, through the degradation of sinking organic
material. Finally, different regions or ecosystems are expected to give
different responses to ocean acidification (Gattuso et al., 2015).
Unfortunately, little has been documented on geographical comparisons of this
aspect.
Relationship of pCO2[seawater] with pH (a) and
dissolved O2(b) in subsurface seawater of the Yellow Sea.
Decline in dissolved O2 combined with an increase in dissolved
inorganic carbon (DIC), as a function of the incubation time (60 h),
in an experiment conducted using subsurface seawater from the East China
Sea (c). Depth ranged from 13 to 75 m for a variety of
subsurface seawater samples, with latitudes at
37∘25′–39∘67′ N and longitudes at
121∘16′–124∘10′ E. Ten 60 mL bottles for dissolved
O2 and ten 60 mL borosilicate glass bottles for DIC wrapped with
black polyethylene were submerged into an in-flow water bath, in which
surface seawater was continuously supplied to control the water bath
temperature. Dark incubated samples were collected after 12, 24, and
60 h of incubation. Seawater samples for the experiment were
collected at 37 m depth on 2 July 2013 using a 10 L Niskin Bottle in the East
China Sea at 28∘50′ N, 122∘15′ E.
Possible forthcoming impacts on ocean acidification
An increase in world population (9 billion estimated in 2050) with
increasing needs of energy, food, medicines, and habitats is one of the key
issues (Mostofa et al., 2013a) that will probably contribute not only to the
increase of atmospheric CO2, but also to the exacerbation of other
factors that may also be related to ocean acidification. Such factors include
enhanced photosynthesis (because of the release of terrestrial OM and
nutrients from increased land use), the increment of OM and nutrients in
wastewater, acid rain, and so on. The following issues can be foreseen in the
next decades, unless remedial actions of some sort are taken:
Long-term homogenous acidification in the deeper waters of both coastal
and oligotrophic oceans, apparently caused by biological respiration of DOM
and PP and their subsequent release of CO2 or DIC, could have key
impacts on marine organisms (Cai et al., 2011; Bates et
al., 2013; Zhai et al., 2013, 2014; Byrne et al., 2010;
Mostofa et al., 2013a). Such homogenous
effects of acidification are directly linked to the effects of GW that can
enhance the surface water temperature. The consequence is an extension of the
summer stratification period, which would determine acidification in deeper
oceans.
Coastal seawater, particularly in locations that are highly influenced
by terrestrial river freshwater inputs, is at risk of substantial
acidification, to a higher extent compared to the open oceans (Zhai et
al., 2014; Thomas et al., 2009; Bate et al., 2013; Barton et al., 2014; Cai
et al., 2011; Cai, 2011; Bauer et al., 2013; Hu et al., 2015). In fact, in
addition to the dissolution of atmospheric CO2, coastal seawater
would be subjected to acidification processes connected with eutrophication,
acid rain, and pollution-affected respiration (Doney et al., 2007; Cai et
al., 2011; Sunda and Cai, 2012; Zeng et al., 2015). Indeed, OM is
substantially increasing in coastal oceans (Bauer et al., 2013). Furthermore,
transport phenomena (e.g., oceanic pump) will gradually increase the level of
nutrients, DOM and PP from coastal areas in the direction of the oligotrophic
open ocean (Fig. 1) (Thomas et al., 2004). Therefore, additional
acidification processes in the oligotrophic open ocean could be operational
and more significant in the coming decades.
Enhanced PP and respiration could increase pCO2 in
open-ocean water and decrease the ability of seawater itself to act as a sink
of atmospheric CO2. The consequence will be an extension of the zones
where seawater acts as a source of CO2, which has increased at an
average rate of 1.5 µatmyr-1 in 1970–2007 (Takahashi et
al., 2002, 2009). In addition to the contribution to ocean acidification, the
decreasing ability of seawater to act as CO2 sink will also
exacerbate the problems related to GW.
The present sea–air fluxes of CO2 (Takahashi et al., 2009)
suggest that the equatorial oceans are prevailingly a CO2 source to
the atmosphere while the temperate ones are mainly a sink. Figure 5 reports
the predicted pH changes by 2100 (Mora et al., 2013), showing that
acidification is expected to affect all the world's oceans but that the most
important effects are predicted for the elevated northern and southern
latitudes. Such locations are presently the sites that mostly act as
CO2 sinks, because seawater pCO2 is lower than the
atmospheric one, and they will experience the most important pH-associated
increase of seawater pCO2. It is thus likely that the global map of
sea–air CO2 fluxes will undergo important changes during the 21st
century.
Possible forthcoming changes in pH in the world's oceans.
Panel (a) shows the spatial difference between future (i.e., the
average from 2091 to 2100) and contemporary (i.e., the average from 1996 to
2005) values under the RCP85 scenario (decadal averages were chosen to
minimize aliasing by interannual variability). Aside each color scale the absolute change is
provided, whereas the numbers on top indicate the
rescaled values; complete results for the RCP85 and RCP45 scenarios for the
ocean surface and floor are shown in the reference (Mora et al., 2013).
Panel (b) shows the global average change relative to contemporary
values under the Representative Concentration Pathways 4.5 (RCP45) and 8.5
(RCP85) scenarios at the ocean surface and seafloor; semi-transparent lines
are the projections for the model.
Impacts of acidification on marine organisms
Marine organisms at low and high latitudes do not respond uniformly to ocean
acidification (Hendriks et al., 2010; Toseland et al., 2013), and the
expected effects can thus be stimulative, inhibitive, or neutral (Anthony et
al., 2008; Gao et al., 2012a; Hutchins et al., 2013). Considering the overall
processes that are involved in ocean acidification (see Fig. 1), it can be
assumed that marine organisms would face detrimental impacts under the
following conditions: (i) they are peculiarly susceptible to pH changes with
different timescales and particularly to acidification, which applies for
instance to the majority of marine calcifiers; (ii) they live under hypoxia
in long-term homogeneous acidified subsurface/deeper seawater, where they
cannot carry out respiration and metabolism properly (this would happen
during a stratification period of increasing duration due to GW, which can
damage their natural growth and development); and (iii) they are subjected to
death/damage in surface seawater by the action of algal toxins and pathogens
(e.g., viruses, coliform bacteria, fungi), and/or to oxidative stress caused
by reactive oxygen species (ROS) and increased water temperature. In many
cases it is extremely difficult (or even next to impossible) to disentangle
acidification from other processes that are taking place at the same time.
Actually, the impacts of increasing acidification on marine organisms may
derive from several processes that are closely interlinked:
(i) acidification, (ii) synergistic effects of acidification and oxidative
stress in surface seawater, (iii) low dissolved O2 (hypoxia) and
acidification in subsurface/deeper seawater, and (iv) stress by algal or
red-tide toxins and pathogens.
Acidification
Impacts induced by seawater acidification or reduced seawater pH are
recognized phenomena and they are discussed in many early reviews. However,
seawater acidification or reduced seawater pH may produce undersaturation of
aragonite and calcite, with the following effects in a variety of seawaters:
(i) dissolution of biogenic shells or skeletons, mostly composed of
CaCO3 in the form of calcite or aragonite, of adult marine
calcifiers such as corals (Kleypas et al., 1999; Erez et al., 2011; Pandolfi
et al., 2011; Wittmann and Pörtner, 2013), crustose coralline algae
(Anthony et al., 2008; Hall-Spencer et al., 2008), shellfish (Talmage and
Gobler, 2010; Barton et al., 2012; Wittmann and Pörtner, 2013), marine
plankton including foraminifera (de Moel et al., 2009; Moy et al., 2009) and
coccolithophores (Riebesell et al., 2000; Beaufort et al., 2011), mollusks
(Doney et al., 2009; Wittmann and Pörtner, 2013), and echinoderms (Doney
et al., 2009; Wittmann and Pörtner, 2013); sedimentary CaCO3
would be affected as well (Kleypas et al., 1999; Bates et al., 2013);
(ii) inability to form new shells or skeletons of framework builders by
larvae or juvenile calcifiers (e.g., the larval and juvenile stages or smaller
individuals), particularly at the early development stages. The effect would
be operational through the decline of calcification rates, which
substantially decreases the growth and development of the organisms including
corals (Kleypas et al., 1999; Anthony et al., 2008; Kroeker et al., 2013);
and (iii) ocean acidification could indirectly enhance heterotrophic
bacterial activities with increasing bacterial protein production and growth
rate at elevated pCO2 levels (Grossart et al., 2006; Endres et
al., 2014; Baragi et al., 2015); higher bacterial abundance has been reported
under high pCO2 treatments (Endres et al., 2014; Tait et
al., 2013), which could consequently accelerate respiration processes and
increase the respiratory CO2 production in the future ocean (Piontek
et al., 2010). As discussed in Sect. 3, seawater pH varies in different timescales and shows short-term variations (e.g., minutes to hours: diurnal and
abrupt) in upper surface seawater and long-term variations (e.g., weeks to
several months: homogeneous) in subsurface and deeper seawater. Long-term
homogenous acidification is apparently responsible for the majority of
impacts on marine organisms. However, the impact on marine calcifiers of pH
variations in different timescales, and most notably the diurnal ones, is
presently poorly known and should be the focus of future research.
Synergistic effects of acidification and oxidative stress in surface seawater
The rapidly rising levels of atmospheric CO2 will result in ocean
warming in addition to lowering the seawater pH (Solomon et al., 2009;
McCulloch et al., 2012). Marine calcifiers are for instance more sensitive to
increased temperature under low pH conditions, because of the combination of
two stressors (Wood et al., 2010; Pandolfi et al., 2011; Hiebenthal et
al., 2013; Kroeker et al., 2013). The synergistic effects of ocean
acidification and oxidative stress, elevated water temperature, or high
irradiance, all connected with increasing CO2 and GW, can affect
marine ecosystems to a variable degree. In some cases the marine primary
productivity is decreased (Boyce et al., 2010; Gao et al., 2012a), while in
other cases the decrease is not so obvious, as tolerance to elevated
CO2 levels may be developed (Feng et al., 2009; Gao et al., 2009;
Connell and Russell, 2010). However, even in the latter instances one may
observe deep changes in species composition (Meron et al., 2011; Witt et
al., 2011), and sometimes even an increase in coral productivity in
experimental studies (Anthony et al., 2008). However, a drop in biodiversity
is generally observed that is always to the detriment of calcifying organisms
(Hall-Spencer et al., 2008; Connell and Russell, 2010). The observed negative
effects include bleaching and productivity loss in coral reef builders
(Hoegh-Guldberg et al., 2007; Anthony et al., 2008), high mortality and
reduction of shell growth and shell breaking force (Hobbs and McDonald, 2010;
Lischka et al., 2011; Hiebenthal et al., 2013), declining calcification and
enhanced dissolution (Rodolfo-Metalpa et al., 2010), decline in abundance of
the juveniles population (Lischka et al., 2011), and increased N : P
ratios of eukaryotic phytoplankton (Toseland et al., 2013).
The mechanism behind the oxidative stress at elevated WT or high irradiance
is caused by a substantial generation of ROS, such as
O2•-, H2O2, HO•, or
1O2, in the surface water layer. The hydroxyl radical
(HO•), a strong oxidizing agent, is produced from either
endogenic or exogenic H2O2 through Fenton and photo-Fenton reactions
in the presence of metal ions, and upon photolysis of NO2- or
NO3- (Zepp et al., 1992; Mostofa et al., 2013c; Gligorovski et
al., 2015). Inside organisms, HO• can damage the photosystem II
activities and finally cause cell death (Blokhina et al., 2003; Mostofa et
al., 2013c). H2O2 concentration levels of approximately
100 nM (compared to up to 1700 nM values that have been detected in
coastal waters) (Mostofa et al., 2013c) can cause oxidative stress to
bacteria, as determined on the basis of increasing catalase enzyme
concentration (Angel et al., 1999). H2O2 can also reduce bacterial
abundances by inducing elevated mortality in seawater (Clark et al., 2008).
The oxidative stress that is related to the Fenton processes would even
increase in acidified water, where the HO• yield is higher
(Zepp et al., 1992). Interestingly and coherently with the expected
HO• yield, the degree of oxidative stress in mollusks has
been found to increase with decreasing pH (Tomanek et al., 2011), and the pH
effect is further exacerbated by an increase in temperature (Matozzo et al.,
2013). Furthermore, the synergistic effect of high H2O2 combined
with high seawater temperature resulted in a 134 % increase in coral
metabolism/respiration rates (Higuchi et al., 2009).
Moreover, one should be sure not to only focus on the direct detrimental effects at the
organism or single-species level: the negative impacts on the dynamics,
structure, composition, and biodiversity of the coral reefs (Findlay et
al., 2010; Wittmann and Pörtner, 2013), of other marine calcifiers (Feng
et al., 2009; Wittmann and Pörtner, 2013), and of marine ecosystem
processes would be linked to changes in species abundance, distribution,
predator vulnerability, and competitive fitness (Hiscock et al., 2004; Feng et
al., 2009; Gao et al., 2012b).
Synergistic effects of low dissolved O2 (hypoxia) and acidification in subsurface/deeper seawater
Declining dissolved O2 in deeper seawater would mostly be caused by
reduced vertical mixing as a consequence of GW (Huisman et al., 2006; Keeling
et al., 2010), which inhibits re-oxygenation while O2 in deep water is
consumed by biological respiration/degradation of sinking organisms and DOM
(Fig. 1) (Stramma et al., 2008; Cai et al., 2011; Sunda and Cai, 2012; Zhai
et al., 2012; Mostofa et al., 2013a). The key reason for hypoxia is the
long-term biological respiration/degradation of sinking OM in the absence of
mixing, which is also a key pathway for acidification in sea subsurface water
during the summer stratification period, as is discussed in earlier sections.
The net decrease of dissolved O2 in subsurface seawater in the Bohai
Sea (China) between June and August 2011 was 34–62 % (see Fig. S3a),
which would be the result of OM respiration during the summer stratification
period. The hypoxia in subsurface water (40 and 70–80 m depths) (Fig. S3b)
along with changes in pH, DOC, and PP or Chl a (Fig. S2)
is linked with enhanced sinking of PP at the end of the summer stratification
period. The connection between hypoxia (through respiration of OM) and
acidification can be assessed by the positive correlation between pH and
dissolved O2 (Fig. S4), which shows that declining O2 is
directly associated with reduced pH in subsurface/deeper seawaters (Fig. S4;
Cai et al., 2011; Zhai et al., 2012, 2013).The connection between hypoxia and
acidification could be exacerbated, and long-term hypoxia could be induced,
by two important factors, namely (i) the increase in algal blooms and the
subsequently enhanced sinking of dead algae in subsurface/deeper seawater,
and (ii) the effects of GW that would induce longer stratification periods as
a consequence of a longer summer season, as previously discussed.
Recent study reveals that hypoxia and acidification have synergistic
detrimental effects on living organisms, because they can separately affect
growth and mortality and their combination can cause damage to organisms
that are resistant to the separate stresses (Gobler et al., 2014). Moreover,
acidification can cause an additional worsening of survival conditions in
oxygen-poor waters, which are already made more acidic by the degradation of
OM (Melzner et al., 2013). The overall consequences of hypoxia and
acidification affect the natural growth and development of organisms (Boyce et
al., 2010) and have implications for habitat loss (Keeling et al., 2010;
Stramma et al., 2010), fish mortality (Hobbs and McDonald, 2010), nutrient
cycling (Keeling et al., 2010; Toseland et al., 2013), carbon cycling
(Keeling et al., 2010), ecosystem functioning (Diaz and Rosenberg, 2008), and
diversity, with possible changes of species composition in the
bentho–pelagic communities (Diaz and Rosenberg, 2008; Stramma et al., 2010).
Stress caused by algal or red-tide toxins and pathogens
Ocean acidification or elevated CO2 could increase the toxic algal
blooms, involving for instance the diazotrophic cyanobacterium
Nodularia spumigena (Endres et al., 2013; Olli et
al., 2015). They could also increase the accumulation of toxic phenolic
compounds across trophic levels in phytoplankton grown under elevated
CO2 concentrations (Jin et al., 2015). Ocean acidification combined
with nutrient limitation or temperature changes could considerably enhance
the toxicity of some harmful groups (Fu et al., 2012). Correspondingly,
harmful algal blooms are expected to increase in coastal waters because of
increasing WT and eutrophication (Anderson et al., 2008; Glibert et
al., 2010; Mostofa et al., 2013a), which would enhance net primary
productivity that is the essential backdrop for the development of such
blooms. The same phenomena are also involved in acidification; thus, it can be
expected that more frequent algal blooms will take place along with ongoing
acidification as an additional stress to marine organisms. Algal blooms and
acidification could also be more closely linked (Cai et al., 2011; Sunda and
Cai, 2012), because the decline of marine algae with a calcareous skeleton
could produce a selective advantage for harmful species (Irigoien et
al., 2005; Mostofa et al., 2013a).
Harmful algal blooms can produce algal toxins (e.g., microcystins) or red-tide
toxins (e.g., brevetoxins) (Flewelling et al., 2005; Anderson et al., 2008),
and the occurrence of pathogens (e.g., potentially hazardous fecal–oral
viruses, coliform bacteria, parasites, or fungi) (Littler and Littler, 1995;
Suttle, 2005) is also more likely in the presence of large phytoplankton
cells and during algal blooms (Fuhrman, 1999; Suttle, 2005). Toxins and
pathogens are a major cause of morbidity and mortality for marine organisms
and they can affect humans as well (Harvell et al., 1999; Flewelling et
al., 2005; Anderson et al., 2008). The most common toxins are microcystins,
cyanotoxins (blue green algal toxins), okadaic acid (OA), dinophysis toxins
(DTXs) and pectenotoxins (PTXs) produced by dinoflagellates (Takahashi et
al., 2007), domoic acid (DA) produced by diatoms (Takahashi et al., 2007),
and brevetoxins produced by the red-tide dinoflagellate
Kareniabrevis (Flewelling et al., 2005; Anderson et al., 2008).
Brevetoxins are potent neurotoxins that kill vast numbers of fish and even
large marine mammals: for instance, 34 endangered Florida manatees
(Trichechus manatuslatirostris) died in southwest Florida in the
spring of 2002, and 107 bottlenose dolphins (Tursiops truncatus)
died in waters off the Florida panhandle in the spring of 2004 as a
consequence of exposure to brevetoxins (Flewelling et al., 2005).
Furthermore, brevetoxins cause illness in humans who ingest contaminated
filter-feeding shellfish or inhale toxic aerosols (Flewelling et al., 2005).
Ocean acidification/elevated CO2 could indirectly affect bacterial
activity and abundance (see Sect. 5.1; Grossart et al., 2006; Allgaier et
al., 2008; Endres et al., 2014; Baragi et al., 2015; Witt et al., 2011; Tait
et al., 2013). However, the abundance of different bacterial communities
could respond differently (increase, remain unchanged or even decrease) under
the effect of global warming (Allgaier et al., 2008; Witt et al., 2011;
Baragi et al., 2015). However, acidification is also connected to an increase
of pathogenic microbiota in corals (Meron et al., 2011). The latter effect is
particularly alarming, because coral reefs are already directly endangered by
acidification (inhibition of the calcification process, as already discussed)
and GW. The reduction in reef-building coral species would be exacerbated by
18 coral diseases identified so far, with increasing prevalence and virulence
in most marine taxa (Sutherland et al., 2004). The most concerning diseases
are the black band disease (BBD), probably caused by several species of
cyanobacteria including most notably Phormidium corallyticum
(Rudnick and Ferrari, 1999); the coralline lethal orange disease (CLOD; a
bacterial disease affecting coralline algae), which greatly impacts coral
reefs and reef-building processes (Rudnick and Ferrari, 1999); a virulent
disease known as white plague type II, which caused widespread mortality in
most Caribbean coral species through physical contact with the macroalga
Halimeda opuntia (Nugues et al., 2004); and, finally, corals
bleaching or disease caused by the temperature-dependent bacteria
Vibrio shiloi (Vidal-Dupiol et al., 2011). Further proposed
pathogens for BBD, in addition to Phormidium corallyticum, include
different genera of cyanobacteria, sulfate-reducing bacteria including
Desulfovibrio spp., sulfide-oxidizing bacteria presumed to be
Beggiatoa spp., several other heterotrophs, and marine fungi (Sekar
et al., 2006). Any bacterial community shifted by elevated CO2 could
thus impact on other marine organisms. Finally, experimental research is
warranted to find out links and mechanisms between harmful algal blooms and
ocean acidification/elevated CO2.
Potential ecological and biogeochemical consequences arising from future ocean acidification
An overview of the potential upcoming ecological and biogeochemical
consequences, linking different environmental drivers, processes, and cycles
related to acidification in the future ocean is provided in Fig. 6. A recent
study demonstrated that different types of tropical cyclones (hurricanes and
typhoons) could increase significantly in oceans and on land over the 21st
century (Lin and Emanuel, 2016). Extreme daily rainfall is thought to
increase with temperature in some regions (Chan et al., 2016 and reference
therein). Watersheds with high precipitation induce higher riverine discharge
rates (Bauer et al., 2013) and, for instance, a single tropical storm can
export approximately 43 % of the average annual riverine DOC (Yoon and
Raymond, 2012). Similarly, on decadal timescales, single large,
cyclone-induced floods can transport 77–92% of particulate organic carbon
from mountainous regions (Hilton et al., 2008). Correspondingly, enhanced
human activities due to increasing population will unquestionably jeopardize
Earth's natural systems. Soil erosion is gradually intensified in regions
where forests are converted into croplands (Ito, 2007), and humans have
increased the sediment transport of global rivers through soil erosion by
2.3±0.6 billion metric tons per year (Syvitski et al., 2005). Potential
changes in erosion rates in the midwestern United States under climate change
is predicted and runoff could increase from +10 to +310 % (along with
soil loss increase from +33 to +274%) in 2040–2059 relative to
1990–1999 (O'Neal et al., 2005). The transfer of OM or organic carbon from
the terrestrial soil to the oceans via erosion and riverine transport could
significantly affect the coastal oceans (Hilton et al., 2008; Bauer et
al., 2013; Galy et al., 2015). Particulate organic carbon
(POC) export from the terrestrial biosphere into the oceans is mostly
controlled by physical erosion, which is thus predicted to become the
dominant long-term atmospheric CO2 sink under a fourfold increase in
global physical erosion rate at constant temperature (Galy et al., 2015).
An overview of the potential upcoming ecological and biogeochemical
consequences, linking different environmental drivers, processes, and cycles
related to acidification in the future ocean.
Such enhanced input of OM with rising temperatures under future global
warming conditions will have a drastic impact on ocean acidification, which is
concomitantly linked with other biogeochemical processes (Jin et al., 2015;
Mora et al., 2013). Moreover, temperature regulates important abiotic and
biotic processes that can alter water throughput, flow paths, dissolution
rates, and watershed carbon stocks (Bauer et al., 2013) as well as
the stratification period or euphotic zone (Fig. 1; Mora et al., 2013; Huisman et
al., 2006; Jöhnk et al., 2008). In addition, elevated temperature under
global warming conditions could potentially enhance the proliferation of
harmful Cyanobacteria in surface water (Paerl and Huisman, 2008; Jöhnk et
al., 2008). The overall ecological and biogeochemical consequences of future
ocean acidification under forthcoming global warming conditions in oceans
could severely impact coastal seas, with a spreading of anoxic dead zones
and a frequent occurrence of toxic dinoflagellate blooms (Jackson, 2008).
Possible evolutions could involve expanding hypoxia in the deeper water
layers (Wannicke et al., 2013; Stramma et al., 2008); changes in food-web
dynamics (Fabry et al., 2008; Wannicke et al., 2013); changes in the
biogeochemical cycling dynamics of C, N, and P (Keeling et al., 2010;
Wannicke et al., 2013; Toseland et al., 2013; Unger et al., 2013; Olli et
al., 2015; Baragi et al., 2015); changes in metabolic pathways (Jin et
al., 2015); increases in coral susceptibility to disease, pathogen abundance,
and pathogen virulence (Maynard et al., 2015); negative consequences to
mortality for various marine organisms, particularly for the shell-forming
ones (Haigh et al., 2015; Doney et al., 2009); structural changes in
phytoplankton communities (Dutkiewicz et al., 2015) and in some marine
keystone species (Waldbusser et al., 2015; Barton et al., 2012); setting up
of the Lilliput effect that causes organisms to evolve towards becoming
smaller and exploit related physiological advantages (Garilli et al., 2015);
increasing appearance of harmful marine species (e.g., Nodularia spumigena sp.; Olli et al., 2015; Jackson, 2008; Paerl and Huisman, 2008)
and of toxic compounds (e.g., of the phenolic type; Jin et al., 2015);
alteration of fish populations through habitat modification (Nagelkerken et
al., 2016), as well as increasing global redistribution of marine
biodiversity (Molinos et al., 2016). Finally, such ecological and
biogeochemical changes in the oceans could have profound consequences for
marine biodiversity, ecosystem services or processes, and seafood quality
with deep implications for fishery industries in the upcoming decades (Doney
et al., 2009; Mora et al., 2013; Jin et al., 2015).
Perspectives
Ocean acidification is the outcome of a series of
anthropic and natural processes that take place at the same time and are
often interlinked. The dissolution of increasing atmospheric CO2 into
seawater obviously plays an important role (Pearson and Palmer, 2000; Feely
et al., 2008; Beaufort et al., 2011), but there are also important
contributions from the degradation of primary producers and DOM (Cai et
al., 2011; Sunda and Cai, 2012; Mostofa et al., 2013a). The latter process
could be enhanced by an increased oceanic primary productivity (Feng et
al., 2009; Sunda and Cai, 2012; Mostofa et al., 2013a), which is one of the
possible consequences of global warming (see also Fig. S1) (Feng et
al., 2009; Mostofa et al., 2013c). In coastal areas, acid rains and
eutrophication caused by the runoff of terrestrial organic matter including
DOM and nutrients (Sunda and Cai, 2012; Bauer et al., 2013), combined with
microbial and photochemical degradation (Mostofa et al., 2013a), may be
important or even the major causes of acidification. All the described
processes would increase the supersaturation of the seawater CO2 that
correspondingly reduces the ability of seawater to take up atmospheric
CO2, thereby extending the oceanic areas that constitute a source
instead of a sink or carbon dioxide (presently, such areas are mostly
concentrated in the equatorial zone) (see Fig. 1). An important issue is that
acidification takes place at varying degrees, with different roles of the
factors involved and with different impacts depending on the latitude, on the
water temperature range as modified by the effects of GW, and on the distance
from the coast (Vitousek et al., 1997; Copin-Montégut et al., 2004; Feely
et al., 2008; Yamamoto-Kawai et al., 2009; Beaufort et al., 2011; Bates et
al., 2013; Kroeker et al., 2013).
Acidification of seawater would be detrimental to marine organisms, and
particularly to marine calcifiers for the long-term (e.g., homogeneous)
acidification of subsurface/deeper seawater and possibly also the short-term
(e.g., diurnal and abrupt) acidification of upper surface seawater. Therefore,
living organisms will have to face multiple stresses at the same time, such
as increasing occurrence of reactive oxygen species in the sea surface water,
hypoxia in subsurface water, toxic algal blooms, and pathogens. Some of these
additional stressors and/or their effects could be enhanced by acidification:
the oxidative stress tends to be more severe at lower pH values and in the
presence of diurnal and abrupt pH variations in surface water; the effects of
hypoxia are exacerbated in long-term homogeneously acidified
subsurface/deeper seawater, and a decline in marine calcifiers could provide
a competitive advantage for toxic algae. Therefore, ocean acidification is
expected to introduce deep changes in marine habitats, and food web
processes.
Based on the discussed mechanisms, some of the possible actions that could be
taken to limit the future impacts of acidification can be listed here
(i) a reduction of anthropic CO2 emissions to the atmosphere, which
should be carried out in the wider context of fighting global warming and
will face the same difficulties; (ii) the implementation of measures aimed at
CO2 capture, such as a worldwide increase in green plantation. In
coastal areas, to limit the effects of acidification, some measures could be
taken that are probably of somewhat easier implementation: (a) reduction of
the inputs to seawater of OM from soil runoff, which implies the control and
limitation of land use practices, of soil erosion and of wastewater
discharges; (b) limitation of the primary productivity by controlling
eutrophication, including the release of nutrients from agricultural
activities; (c) removal of algae (e.g., by means of nets) during bloom
periods, to avoid fertilization of seawater by the associated nutrients;
(d) limitation of the emission of pollutants such as nitrogen and sulfur
oxides to the atmosphere, as they are precursors of HNO3 and
H2SO4 that are involved in acid rains. Finally, marine
oceanographers should focus on how marine organisms are affected by
short-term pH variations (e.g., diurnal and abrupt) in surface waters and by
long-term (e.g., homogeneous) ones in response to the effects of GW, which may
further influence such pH variations.
The Supplement related to this article is available online at doi:10.5194/bg-13-1767-2016-supplement.
Acknowledgements
This work was supported by the National Natural Science Foundation of China
(grant nos. 41210004, 41130536 and 41276061). MM and DV acknowledge financial
support by University of Torino – EU Accelerating Grants, project
TO_Call2_ 2012_0047 (Impact of radiation on the dynamics of
dissolved organic matter in aquatic ecosystems – DOMNAMICS). This study was
also partly supported by the Grants-in-Aid for Scientific Research
(no. 09041159) for International Geosphere–Biosphere Programme (IGBP) at
Nagoya University from the Ministry of Education, Culture, Sports, Science
and Technology (MEXT) of Japan and also by a grant from the LBRI and partly a
Grant-in-Aid for the IHP fellowship awarded to Khan M. G. Mostofa from the
MEXT.
Edited by: G. Herndl
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