Post on 20-May-2020
Chapter 31
Ocean Circulation and Climate, Vol. 103. http://dx.doi.org/10.1016/B978-0-12-391851-2.00031-3
Copyright © 2013 Elsevier Ltd. All rights reserved.
Marine Ecosystems, Biogeochemistry,and Climate
Scott C. DoneyWoods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA
Chapter Outline1. Introduction 817
2. Phytoplankton, Primary Production, and Climate 820
3. Climate Impacts on Higher Trophic Levels 824
4. Ocean Acidification 828
5. Deoxygenation and Hypoxia 830
6. Marine Biogeochemical Cycles–Climate Interactions 831
7. Observational and Research Directions 833
Acknowledgments 834
References 834
1. INTRODUCTION
This chapter introduces key basic concepts on marine eco-
system dynamics, discusses how physical variability
impacts ocean biota on large scales (i.e., gyre to basin, inter-
annual to centennial), and touches on how ocean biogeo-
chemical processes can modify physical climate
(Sarmiento and Gruber, 2006; Miller and Wheeler, 2012).
An ecosystem consists of a complete set of biotic and
abiotic components of a system or location—all the living
organisms, nutrients and detrital materials, and the physical
environment—as well as the interactions among all of these
components. For the pelagic ocean, key organism groups
span from microscopic phytoplankton and bacteria through
zooplankton all the way up the trophic ladder to fish, marine
mammals, and seabirds (Figure 31.1). Relevant biological
interactions include inter-species competition, predator–
prey relationships, disease, and parasitism. Important
physical processes involve seawater chemistry, temperature
and light, vertical and horizontal turbulent mixing, and
ocean circulation that helps govern nutrient supply and
the dispersal of organisms. The spatial extent of an eco-
system is defined more by the strength of the interactions
rather than by spatial homogeneity. Of course, no marine
ecosystem is fully self-contained, and constraining physical
and biological transport (Williams and Follows, 2011) often
is essential to understand the functioning of the ecosystem.
Given the focus here on ecosystem–climate scale inter-
actions, by necessity, the chapter neglects the many
fascinating smaller-scale biological–physical phe-
nomenon—for example, microscale diffusion and
molecular viscosity; local-scale Langmuir cells and internal
waves; and mesoscale fronts and eddies (Mann and Lazier,
2005)—that are critical for maintaining the base state upon
which climate variability acts. Similarly, one cannot hope to
capture in a single chapter, a detailed discussion on the hier-
archy of biological scales from individual organisms and
populations of distinct species up through the interacting
communities of different species that compose the core of
an ecosystem. Nor, in many cases, do we have sufficient
information to document climate–biological interactions
across all biological scales, particularly at the community
and ecosystem level. Rather, current understanding depends
heavily on theory and laboratory experimental results at the
organism level to help explain observed spatial patterns and
historical trends in aggregated measures such as phyto-
plankton primary production or the abundance and spatial
range of particular species. The emphasis in the chapter
falls particularly on understanding the biological response
to observed physical variability and trends over the twen-
tieth and early twenty-first centuries as well as projecting
the potential impacts on marine ecosystems due to further
anthropogenic climate change over the next several decades
to centuries.
Marine ecosystems are already experiencing large-scale
trends in physical climate, ocean chemistry, and other
human environmental perturbations (Figure 31.2) (Doney,
817
FIGURE 31.1 Food-web schematic for the coastal waters
near Palmer Station on the West Antarctica Peninsula. Sig-
nificant regime shifts are occurring in the marine ecosystem
in this region linked to warming and sea-ice decline, and
decade time-scale shifts in population levels are confirmed
from field observations for the biological groups indicated
by the red dashed box (Ducklow et al., 2007; Schofield
et al., 2010). Figure courtesy of Hugh Ducklow.
PART VI The Changing Ocean818
2010; Gruber, 2011; Doney et al., 2012). Documented
physical climate changes relevant to marine biota include
rising sea-surface temperature (SST), upper-ocean
warming, sea-level rise, altered precipitation patterns and
river runoff rates, and sea-ice retreat in the Arctic and west
Antarctic Peninsula (Figure 31.2) (Bindoff et al., 2007).
Reduced stratospheric ozone over Antarctica appears
to be causing a major shift in atmospheric pressure
(more positive Southern Annular Mode conditions), which
strengthens and displaces poleward the westerly winds in
the Southern Ocean and which also may be increasing
ocean vertical upwelling. Future climate projections
indicate continuation and, in many cases, acceleration of
these trends as well as other changes such as more intense
Atlantic hurricanes (Bender et al., 2010), an ice-free
summer in the Arctic (Stroeve et al., 2012), and a very likely
reduction in the strength of the Atlantic deepwater for-
mation (Bryan et al., 2006).
Relevant chemical trends include rising seawater CO2
levels (leading to ocean acidification) (Gattuso and
Hansson, 2011), reduced dissolved oxygen (O2) concentra-
tions reflecting warming and altered circulation (deoxygen-
ation) (Keeling et al., 2010), and growing coastal nutrient
levels leading to eutrophication, and expanding coastal
and estuarine hypoxia (very low dissolved O2) (Rabalais
et al., 2010). These chemical trends are caused by the same
global human pressures driving climate change, namely,
fossil-fuel burning, deforestation, and industrial-scale agri-
culture (Le Quere et al., 2009). Climate change also may
exacerbate the ecosystem impacts of other human pressures
and stressors such as coastal habitat loss, coastal urbani-
zation, and overfishing, which have increased in magnitude
dramatically over the past several decades (bottom panel
Figure 31.2; Doney et al., 2012). Therefore, organisms
(and ecosystems) will experience simultaneously multiple
physical and chemical stressors that may exceed their capa-
bility to acclimate or adapt (Boyd et al., 2008).
The physical environment directly influences
organism physiology through multiple pathways including
temperature, salinity, O2, CO2, pH, etc. (Somero, 2012).
Temperature variations and thermal stress are perhaps
the most straightforward to understand because most
marine plants and animals are ectothermic and cannot
regulate their internal body temperature. Therefore, sea-
water temperature plays a central role modulating
almost all biological rates, including growth and repro-
duction as well as microbial processes that dominate
ocean biogeochemical cycling. Metabolic rates tend to
rise exponentially with temperature up to some threshold
temperature, above which thermal stress kicks in and bio-
logical rates drop sharply (Portner and Farrell, 2008). The
exponential relationship with temperature can be captured
by a Q10 value, that is, the rate increase resulting from a
10 �C rise in temperature. For example, Eppley (1972)
reported a Q10 of �1.9 for the upper envelope of growth
rates among �130 species and clones of phytoplankton;
a 2 �C warming would, therefore, yield a 37% increase
in growth rate.
1900 1920 1940 1960 1980 2000 2020250
300
350
400
Year
Physical climate forcing
Atm
osph
eric
CO
2 (p
pm)
−1
−0.5
0
0.5
SS
T a
nom
aly
(�C
)
1900 1920 1940 1960 1980 2000 20200
0.2
0.4
0.6
0.8
1
1.2
Year
Rel
ativ
e to
yea
r 20
00
Other human perturbations and impacts
U.S. coastal populationAnthropogenic nitrogen fixationMarine wild-fish harvestCumulative hypoxic zonesCumulative Caribbean coral cover lossCumulative seagrass lossCumulative mangrove loss
FIGURE 31.2 Time-series trends over the twentieth and early twenty-first centuries for physical climate and anthropogenic perturbations relevant to
marine ecosystem dynamics. Top panel: annual-average atmospheric CO2 from ice cores prior to 1959 (MacFarling Meure et al., 2006) and Mauna Loa
instrumental record from 1959 to present (Tans and Keeling, 2012); and global-mean SST anomalies (ERSST data referenced to 1971–2000 climatology)
(Smith et al., 2008). Lower panel: U.S. coastal population (Wilson and Fischetti, 2010), anthropogenic nitrogen fixation (Davidson, 2009), global marine
wild-fish harvest (Food Agric. Org. U.N., 2010), cumulative global hypoxic zones (Diaz and Rosenberg, 2008), cumulative seagrass loss (Waycott et al.,
2009), cumulative Caribbean coral cover loss (Gardner et al., 2003), cumulative mangrove loss (Food Agric. Org. U.N., 2007). All time series in lower
panel are normalized to 2000 levels. Adapted from Doney et al. (2012).
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 819
On this basis, it might be expected that primary pro-
duction, as well as the growth rates of ectothermic animals
and pathogens, will increase in a warmer ocean. However,
nutritional status, thermal tolerance, O2 availability, envi-
ronmental chemistry, food availability, or other factors
may limit growth and production or other biological pro-
cesses, regardless of metabolic rate. Further, most
organisms inhabit a geographic range often bounded by
upper and lower temperature limits, which may be further
constrained by biological interactions with prey, compet-
itors, predators, parasites, and diseases. As climate warms,
species’s geographic ranges may shift poleward to maintain
a similar thermal niche, all other factors remaining
favorable.
Over evolutionary time scales, organism life histories
adapt to the physical climate and biological community
in which the species population is embedded. Rapid envi-
ronmental variability on short time scales can disrupt key
biological relationships that underpin an organism’s food
supply or reproductive success. Many species exhibit
seasonal variations in the timing or phenology of major life
events such as reproduction. Changes in the spatial pattern,
abundance or timing of prey blooms, for example, could
result in dramatic indirect climate impacts on a predator.
Climate variations and trends can create mismatches in time
or space due to differential responses of species, potentially
leading to cascading effects through a food web (Edwards
and Richardson, 2004; Parmesan, 2006). For example, the
seasonal match/mismatch in the timing of fish larval pro-
duction to planktonic food supply has been suggested as
an important factor driving year to year variability in fish
recruitment (e.g., Cushing, 1990). This could translate into
substantial and nonlinear biological responses to climate
change from shifts in phytoplankton and zooplankton phe-
nology (Stenseth and Mysterud, 2002). Organisms attempt
to cope with such disruptions through physiological accli-
mation, behavior modifications, and eventually evolu-
tionary adaption. The cumulative direct and indirect
climate responses of individual organisms and species
populations alter aggregated properties of an ecosystem
PART VI The Changing Ocean820
such as primary production, energy and mass flow, com-
munity structure, and biodiversity.
The remainder of the chapter is organized as follows.
Section 2 discusses the influence of physics and climate
variability on phytoplankton distributions and primary pro-
duction. This is followed by a survey of climate impacts on
higher trophic levels, focusing primarily on thermal effects
that are relatively more well documented in the literature
compared to most other stressors (Section 3). Section 4
touches on the seawater chemistry changes associated with
rising atmospheric CO2 as well as the biological responses
to the resulting ocean acidification. Section 5 highlights the
effects of climate and nutrient eutrophication on ocean O2
distributions. Section 6 talks about the coupling between
marine biogeochemistry and global climate in terms of
ocean CO2 storage and O2 distributions as well as
climate-active trace gases. The chapter concludes with a
brief discussion on future observational and research direc-
tions (Section 7). The chapter draws on and builds from
several recent review articles on carbon cycle-climate cou-
pling (Doney and Schimel, 2007), ocean acidification
(Doney et al., 2009), ocean biogeochemistry (Doney,
2010), and climate change impacts on ocean ecosystems
(Doney et al., 2012; Griffis and Howard, 2012).
2. PHYTOPLANKTON, PRIMARYPRODUCTION, AND CLIMATE
In the upper-ocean, small floating photosynthetic microbes
and plants, collectively called autotrophic phytoplankton,
use sunlight to convert inorganic CO2 into organic matter
and O2 via the simplified net overall equation for
photosynthesis:
CO2þH2O)CH2OþO2 ð31:1Þwhere CH2O is a generic carbohydrate. Associatedmetabolic
processes, such as synthesis of proteins and enzymes, DNA
and RNA, and lipids, also require bioavailable forms of
nitrogen, phosphorus, and trace elements, most notably iron
(Geider et al., 1997). Phytoplankton growth rates are gov-
erned “bottom-up” by temperature, light, and limiting macro
and micronutrients. Typically, growth rates increase linearly
with light or nutrients at low illumination and nutrient levels,
eventually saturating at a temperature-dependent maximum
growth rate. Diatoms also require silicon to build their shells,
whereas coccolithophores need carbonate ions (CO2�3 ) to
build calcium carbonate (CaCO3) shells. The local time rate
of change in phytoplankton biomass (P) depends on physicaladvection, mixing, sinking, and the net balance of biological
growth and loss terms:
@P
@tþr� u⇀Pð Þ�r� KrPð Þ¼RHSbio ð31:2Þ
where u⇀ is velocity and K is turbulent diffusivity. The bio-
logical right-hand-side terms, RHSbio, can be expressed as
the net specific growth rate m:
1
P
dP
dt¼ m¼ photosynthesis�grazing�other loss terms
ð31:3Þwhere “top-down” losses are dominated by zooplankton
grazing as well as other, less well quantified, processes such
as viral lysis, cell death, and phytoplankton aggregation that
leads to gravitational sinking out of the well-lit upper ocean.
The stored chemical energy from phytoplankton
primary production supports rich pelagic food webs in both
the coastal and open-ocean, including deep sea and benthic
ecosystems (Figure 31.3). Recent estimates for globally
integrated marine net primary production are in the range
of 60–70 PgC year�1 (where 1 Pg¼1015 g) (Behrenfeld
et al., 2005). Most of the organic carbon produced by phy-
toplankton is converted back to CO2 in the upper ocean
through respiration (the reverse of Equation 31.1). The
primary loss mechanisms are via cycling through the het-
erotrophic bacterial loop or grazing by zooplankton.
Transfer of organic carbon to higher trophic levels—that
is fish, marine mammals, etc.—is inefficient, and marine
biogeochemical mass and energy cycling are dominated
by the activity of microbes and plankton. A small and rel-
atively uncertain fraction of primary production is exported
into the subsurface ocean, roughly 5–12 PgC year�1
(Dunne et al., 2007; Henson et al., 2011), where respiration
releases CO2 and nutrients and consumes O2. Export flux is
modulated by phytoplankton size structure with a greater
fraction of production export from regions with larger cells
and especially diatoms with siliceous shells. As a result of
the net fixation of organic carbon in the euphotic zone and
respiration in deeper waters, surface waters tend to have
lower dissolved inorganic carbon (DIC) levels, whereas
thermocline and deep waters are marked by higher DIC
and nutrient concentrations and lower O2, even after
accounting for variations in temperature-dependent solu-
bility (cold water holds more gases than warm water).
Relative to terrestrial systems, the elemental stoichi-
ometry of marine plankton and sinking particles is rela-
tively uniform, a fact first noted in a series of seminal
papers (Redfield, 1958; Redfield et al., 1963) and codified
in the so-called Redfield ratios relating the molar ratio
of P:N:C:O2 during net production and remineralization
(1:16:117: �170) (Anderson and Sarmiento, 1994). Red-
field ratios provide conversion factors for interrelating
different types of ocean biogeochemical measurements of
new production, net community production, and export
flux that often are constructed from different elemental
currencies. Although fixed Redfield ratios are a good guide,
recent work indicates systematic spatial and temporal var-
iations in plankton elemental composition in response to
FIGURE 31.3 Schematic of the flow of organic carbon through a generic open–ocean pelagic food web, the so-called biological carbon pump that trans-
ports carbon from the surface ocean to the deep sea and increases natural ocean carbon storage (left panel). A schematic of the physiochemical “solubility”
carbon pump driven by CO2 solubility and ocean circulation is also shown (right panel). The thicknesses of the light blue bands in the left panel indicate that
most of the organic carbon produced by phytoplankton primary production is respired in the upper ocean by bacteria, zooplankton, and, to a smaller degree,
higher tropic levels. Export of organic carbon to the deep sea typically is a small fraction of primary production. Figure from Chisholm (2000).
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 821
variations in community structure, nutrient stress, and other
biological factors (Geider and LaRoche, 2002; Deutsch and
Weber, 2012; Martiny et al., 2013).
The large-scale patterns of phytoplankton biomass,
mapped from satellite remote sensing in terms of concen-
tration of the photosynthetic pigment chlorophyll (ocean
color), broadly reflect the patterns of the wind-driven gyre
circulation and coastal upwelling (McClain, 2009; Siegel
et al., 2013; Figure 31.4). Export of sinking organic matter
strips surface waters of vital nutrients over time, in effect-
transporting nutrients diapycnally from light surface waters
to dense thermocline and deep waters. Biomass levels,
therefore, are modulated by physical processes that result
in upward fluxes of nutrient-rich subsurface waters—
large-scale upwelling, seasonal convection, and mesoscale
eddies (McGillicuddy et al., 2003). Surface chlorophyll
levels are low in nutrient-poor, subtropical gyres charac-
terized by downwelling and deep thermoclines. In contrast,
surface chlorophyll levels are more than an order of mag-
nitude higher in nutrient-rich subpolar waters marked by
large-scale upwelling of cold, nutrient-rich water and
shallow thermoclines. Chlorophyll also can be considerably
greater on continental shelves and upwelling eastern
boundary current systems such as the California Current
off the west coast of North America and the Benguela
Current off the west coast of southern Africa. The ultimate
source of thermocline nutrients and, thus, low-latitude
productivity involves global-scale circulation; wind-driven
upwelling in the Southern Ocean forms nutrient-rich mode
and intermediate waters that then flow northward at mid-
depth indirectly supplying the nutrients that feed produc-
tivity over much of the globe (Sarmiento et al., 2004a).
Some surface ocean regions have abundant macronu-
trients but relatively low chlorophyll, and phytoplankton
growth appears limited by surface iron levels (Martin and
Fitzwater, 1988; Boyd et al., 2007). Sources of iron to the
surface ocean include atmospheric dust deposition, conti-
nental shelf sediments, and upwelling of recycled iron
(Jickells et al., 2005). Inputs of bioavailable iron are rela-
tively low in the subpolar North Pacific, eastern Equatorial
Pacific, and the Southern Ocean, resulting in High-Nitrate,
Low Chlorophyll (HNLC) conditions, under which iron
limitation slows phytoplankton growth rates, particularly
for larger cells, and reduces export fluxes.
Spatial and temporal variations in surface solar radiation
and mixed layer depth also affect phytoplankton pro-
duction. A little less than half of total solar irradiance falls
in the bands classified as photosynthetically available
radiation (PAR) (�400–700 nm), and PAR levels drop
approximately exponentially with depth away from the
ocean surface. Deeper mixed layers, therefore, reduce the
average light level seen by the phytoplankton community
over the mixed layer and, thus, lower the average photo-
synthesis rate. Large seasonal phytoplankton blooms occur
FIGURE 31.4 Satellite estimated ocean surface chlorophyll concentration from the Sea-viewingWide Field-of-view Sensor (SeaWiFS) using the OC4v6
band-ratio chlorophyll algorithm. Mean values over the length of the SeaWiFS mission (1997–2010) are calculated for 1� bins in latitude and longitude
over the global ocean. Units are log10(mg m�3). The mean SST¼15 �C isotherm is shown as the black lines. Figure from Siegel et al. (2013).
PART VI The Changing Ocean822
in late-winter through spring and early summer in many
temperate waters; the traditional explanation in terms of
the Sverdrup’s Critical Depth Hypothesis suggests that
blooms are triggered by the relief of community light lim-
itation due to increasing surface irradiance and shoaling
mixed layers (Siegel et al., 2002). Alternatively, bloom ini-
tiation can be traced back to mixed layer deepening earlier
in the season, which decouples the zooplankton–phyto-
plankton grazing relationship that normally keeps net phy-
toplankton growth m near zero (Evans and Parslow, 1985;
Behrenfeld, 2010). Reconciliation of these two hypotheses
may lie in recognizing that community respiration varies
over time and that grazing and light-driven variations in
productivity more strongly influence bloom dynamics
during different stages of the seasonal cycle.
Physical and biological factors influence plankton com-
munity composition as well as primary production. For
example, low nutrients in the subtropical oligotrophic ocean
favor species with smaller cells, prokaryotic and small
eukaryotic picoplankton, of the order of O(1)mm in
diameter, and nanoplankton of the order of O(10)mm in
diameter. Cells adapt to severe phosphorus limitation by
replacing phosphorus-based lipids in cell membranes with
unique sulfur and nitrogen-based lipids (Van Mooy et al.,
2009). Nitrogen-fixing diazotrophic organisms, which can
create bioavailable nitrogen from otherwise inert N2 gas,
also arise in low-nitrate oligotrophic waters (Karl et al.,
1997). Warm, well-stratified oligotrophic waters, typically,
are characterized by a microbial food web with low
biomass, rapid recycling of organic matter by small micro-
zooplankton, long complex food chains, reduced export
flux and elevated production of dissolved organic matter
and bacterial activity. At the other extreme, nutrient-rich,
productive, coastal, and polar waters typically contain both
small and large phytoplankton cells and exhibit higher
export rates. The larger cells are often dominated by bloom
forming diatoms (10–200 mm in diameter) grazed by larger
meso- and macro-zooplankton that, in turn, support rela-
tively shorter, more direct food chains to higher trophic
level predators such as fish.
Associated with El Nino-Southern Oscillation (ENSO)
and other climate modes, marine phytoplankton and
primary production exhibit substantial climate-driven inter-
annual variability (Henson et al., 2009), estimated, on a
global scale from satellite remote sensing, to be roughly
�2PgC year�1 (Chavez et al., 2011), that is, a few percent
of the total; regional variations can be substantially larger in
a fractional sense. Satellite ocean color records for the past
couple of decades indicate a robust anticorrelation of
tropical and subtropical surface chlorophyll to SST,
upper-ocean heat content, and thermocline depth
(Figure 31.5; Behrenfeld et al., 2006). For example, one
of the largest interannual signals in satellite ocean color
and in situ chlorophyll involves a shift from low to high
surface chlorophyll in the tropics and subtropics linked to
the transition fromwarmEl Nino to cold La Nina conditions
in 1998–1999 in the tropical Pacific (Chavez et al., 1999;
McClain, 2009). These findings are consistent with argu-
ments that increased vertical stratification limits nutrient
supply in stratified waters. However, recent work (Siegel
et al., 2013) suggests a more subtle interpretation of the sat-
ellite ocean color record that most of the chlorophyll vari-
ability in the tropics and subtropics reflects physiological
adjustments in intracellular chlorophyll concentrations
rather than biomass variations; nutrient supply could still
be the distal cause of the observed chlorophyll variations,
but the signal is then more one of change in cell health
rather than abundance. Ocean color variations in temperate
and high latitudes appear to be driven more by changes in
phytoplankton biomass; the northern hemisphere waters
FIGURE 31.5 Time series of SeaWiFS surface ocean chlorophyll (Chl)
(Figure 31.4) and sea-surface temperature (SST) monthly standardized
anomalies (z-scores) for three global regions delineated by the mean
SST isotherm for (a) the cool (mean SST<15 �C) northern hemisphere
(NH) aggregate, (b) the warm, permanently stratified ocean aggregate
(mean SST>15 �C) and (c) the cool southern hemisphere (SH) (mean
SST<15 �C). Anomalies are constructed by first removing the monthly
mean value for each 1� bin of each property and then aggregating the
regional, monthly anomalies into global aggregates. Figure from Siegelet al. (2013).
ΔPP (mgC m-2 day-1)
�400 �100 �25 0 25 100 400
FIGURE 31.6 Projected climate-driven changes in vertically integrated,
annual mean net primary production by the end of the twenty-first century
(difference between 2090–2099 and 1860–1869 decadal means). Multi-
model means are weighted by model skill under contemporary conditions,
and dotted areas indicate regions where all of the models have low skill
scores. Figure from Steinacher et al. (2010).
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 823
also exhibit an anticorrelation between chlorophyll and
SST, with a less clear relationship in the southern hemi-
sphere. Climate-driven variability in surface chlorophyll
and primary production is also abundantly evident in
longer, multidecade in situ time series such as BATS,
HOT and more coastal stations (Chavez et al., 2011).
Discerning decadal and longer-term trends in phyto-
plankton and primary production from existing in situand satellite observations is more difficult because of the
short time series available and the presence of substantial
natural interannual variability (Saba et al., 2010; Chavez
et al., 2011). Modeling studies suggest some caution in
the interpretation of historical data, arguing that it may
require several decades or more of observations to clearly
discern any anthropogenic signal (Henson et al., 2010;
Beaulieu et al., 2013). Boyce et al. (2010) published a pro-
vocative result indicating an �1% year�1 decline in global
median chlorophyll over the past century based on ocean
transparency and in situ chlorophyll data; this would
indicate a wholesale change in ocean circulation and marine
ecosystems. Other researchers, however, argue that the
decline is an artifact of temporal sampling bias or merging
of different data types (e.g., Rykaczewski and Dunne,
2011). Bridging satellite ocean color data between the
Coastal Zone Color Scanner (CZCS; 1979–1986) and initial
SeaWiFS data (1998–2002), Antoine et al. (2005) found a
22% increase in global average chlorophyll. Using
SeaWiFS data, Polovina et al. (2008) observed a 15%
increase over 1998–2006 in the spatial extent of the most
oligotrophic surface waters (�0.07 mg Chl m�3), but this
time-slice includes the large 1997–1998 El Nino event
and subsequent La Nina (Figure 31.5).
Anthropogenic climate change impacts on phyto-
plankton are expected to grow with time over the twenty-
first century in response to further upper-ocean warming
and increased vertical stratification. The resulting decline
in nutrient supply into subtropical surface waters is pro-
jected to reduce primary production (Figure 31.6;
Sarmiento et al., 2004b; Doney, 2006; Steinacher et al.,
2010) and increase nitrogen fixation (Boyd and Doney,
PART VI The Changing Ocean824
2002). The situation is less clear in temperate and polar
waters though there is a tendency in most models for
increased production due to warming, reduced vertical
mixing, and reduced sea-ice cover (Bopp et al., 2013).
The spatial extent of biomes may expand or contract with
the potential for the emergence of new regions with com-
bined biotic and abiotic conditions that have not been
observed before. For example, Polovina et al. (2011)
forecast an �30% expansion in the spatial extent of the
North Pacific subtropical biome by 2100 due to stratifi-
cation and a poleward shift of the mid-latitude westerly
winds. Also, a novel thermal habitat is created in the area
of very warm tropical and subtropical surface waters, mean
annual SST exceeding 31 �C, growing in the simulation
from a negligible amount to over 25 million km2. Higher
SSTs likely will cause poleward migration of phytoplankton
thermal niches and may sharply reduce tropical phyto-
plankton diversity (Thomas et al., 2012). Warming also
may cause the fraction of small phytoplankton (i.e., picophy-
toplankton) to increase, reducing the energy flow to higher
trophic levels (Moran et al., 2010; Marinov et al., 2010).
Marine phytoplankton also can influence ocean physics
and climate. In the upper ocean, the vertical attenuation of
solar radiation depends strongly on the abundance of chlo-
rophyll as well as biologically derived detritus and colored
dissolved organic matter (Morel and Antoine, 1994). In
more productive regions, solar radiation is absorbed closer
to the surface, resulting in shallower mixed layers, warmer
SSTs, and altered air–sea heat and freshwater fluxes, partic-
ularly in the tropics and subtropics (Ohlmann et al., 1996).
Nonlocal and often nonintuitive effects can arise because
altered upper-ocean stratification affects ocean currents
and heat transport (Sweeney et al., 2005), and modeling
studies indicate that ocean chlorophyll may substantially
influence equatorial Pacific Ocean thermal structure
(Murtugudde et al., 2002), modify ENSO interannual vari-
ability (Jochum et al., 2010), and steer tropical cyclones
(Gnanadesikan et al., 2010).
3. CLIMATE IMPACTS ON HIGHERTROPHIC LEVELS
Climate-driven variations in primary production introduce
bottom-up effects on ocean food webs that couple with direct
impacts on higher trophic level organisms. These include
physiological intolerance to changing physical and chemical
environments, altered dispersal and migration patterns, and
shifts in species interactions such as predation and compe-
tition. Population-level changes may occur in abundance,
spatial range, and seasonal timing of major species’s life
history events or phenology. Populations of different
organisms interacting through predator–prey, competition,
and parasite–host relationships constitute a biological
community. Together with local climate-driven invasion
and local (and perhaps global-scale) extinction, climate pro-
cesses may result in altered community structure and
diversity, including possible emergence of novel ecosystems.
Biological variability at a particular trophic level can be clas-
sified as controlled resources (bottom-up) or predation (top-
down); in wasp-waist ecosystems, population variations in a
crucial intermediate trophic level can generate both top-down
influences on lower trophic levels and bottom-up influences
on higher predators (Cury et al., 2000).
Ocean biological populations exhibit substantial inter-
annual to decadal variability, and in many cases, a sub-
stantial component of the variability can be correlated
with global-scale climate modes such as ENSO, the Pacific
Decadal Oscillation (PDO), or the North Atlantic Oscil-
lation (NAO) (Stenseth et al., 2003). Climate-related vari-
ability has been identified across a wide range of
biological taxa from zooplankton (Brodeur and Ware,
1992; Mackas and Beaugrand, 2010) to commercial fish
species such as North Pacific salmon and groundfish
(Mantua et al., 1997; Hollowed et al., 2001).
Zooplankton cover a wide range of taxonomic groups
and play a pivotal role in marine food webs because they
feed directly on phytoplankton, bacteria, and often other
smaller zooplankton. Zooplankton, in turn, serve as prey
for many fish, marine mammals, and seabirds. Long records
of zooplankton abundance and community composition are
available for the eastern North Atlantic from the Con-
tinuous Plankton Recorder (CPR) survey; the CPR data
exhibit climate-related trends and variability in, for
example, calanoid copepod (zooplankton crustaceans) bio-
diversity and biogeographic distributions (Beaugrand et al.,
2002) associated with warming and shifts in the NAO.
At subtropical open-ocean time series sites off Hawaii
and Bermuda, mesozooplankton (>200 mm) dry-weight
biomass is increasing with time and is positively correlated
with SST at Bermuda (Steinberg et al., 2012). These sites
exhibit positive or neutral primary productivity trends
(Saba et al., 2010), despite surface warming that is expected
to decrease both primary and secondary production. Pos-
sible explanations for the zooplankton trends include a
combination of transient responses, increased nitrogen fix-
ation supporting great primary production, alteration in top-
down controls, and northward expansion of tropical
species’ ranges.
Other interesting zooplankton–climate relationships
arise from the California Cooperative Oceanic Fisheries
Investigations (CALCOFI) data in the California Current
System. The CALCOFI data exhibit, among other features,
increases in cold-water krill species on interannual scales in
association with strong coastal upwelling and La Nina
events as well as decadal variations in the abundance of
warm water, coastal krill reflecting changes in the strength
of poleward transport (Brinton and Townsend, 2003;
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 825
Bestelmeyer et al., 2011). The CALCOFI data also doc-
ument long-term warming and declining zooplankton
biomass trends (Roemmich and McGowan, 1995), pri-
marily reflecting decreasing abundance of pelagic tunicates
(Lavaniegos and Ohman, 2003), possibly because warm-
water taxa (small subtropical copepods) were favored by
a shift toward greater relative abundance of small prey
(picoplankton). Climate-driven fluctuations in individual
zooplankton populations in the California Current system
also shift planktonic food-web structure and dynamics
(Francis et al., 2012).
Boundary current systems exhibit striking decadal-scale
oscillations in the abundance of small pelagic fish
(Kawasaki, 1992). In the Pacific, oscillations between
anchovies and sardines are approximately synchronous in
the California Current and Peru/Chile (Humboldt Current)
upwelling systems as well as in the Kuroshio, with
anchovies dominating during cold, negative phases of the
PDO, replaced by sardines during warm phases (Chavez
et al., 2003). The large-scale synchrony must reflect atmo-
spheric teleconnections, though the exact mechanisms
are not fully resolved (Alheit and Bakun, 2010).
Rykaczewski and Checkley (2008), for example, suggest
that the abundance of Pacific sardines covaries with
wind-stress curl-driven upwelling that has relatively low
vertical velocity compared to near-shore coastal upwelling
caused by along-shore winds; they hypothesize that sar-
dines are favored because of the smaller-sized plankton
assemblage under those conditions.
Marine ecosystem structure as a whole also appears to
vary at decadal time scales with some of the best explored
examples involving the response of the North Pacific eco-
system to the PDO. Hare and Mantua (2000) and others
argue that the system undergoes periodic regime shifts, that
is, relatively abrupt, synchronous transitions from one
quasi-stable state to another on decadal time scales in con-
junction with changes between cold and warm phases of the
PDO. Ocean dynamics acts to integrate high-frequency
atmospheric weather resulting in low-frequency (red-
spectrum) variance of ocean properties that could appear
at first glance like a regime shift, but which, in actuality,
is simply a linear response to external forcing (Doney
and Sailley, 2013). Thus, it is often difficult to determine
whether decadal ecological trends simply track external
physical forcing or reflect more complicated, nonlinear
responses that may arise from internal biological interac-
tions (Overland et al., 2010). Careful analysis suggests that,
at least to some degree, the North Pacific regime shift is real
and reflects a combination of forced linear responses and
internal nonlinear biological dynamics (Hsieh et al.,
2005). Nonlinear regime shifts have been detected in
several other marine population time series suggesting
the potential for more climate-driven ecological surprises
(Bestelmeyer et al., 2011).
Turning to anthropogenic climate change, polar marine
ecosystems appear likely to be particularly sensitive to
climate-driven sea-ice variations, which can substantially
restructure food-web pathways from plankton through to
higher trophic levels (Ducklow et al., 2007; Grebmeier,
2012). For example, along the western side of the Antarctic
Peninsula seasonal sea-ice duration has declined by nearly
90 days since the beginning of satellite-based measure-
ments in 1978, reflecting both warming of the ocean-
atmosphere system and strengthening of wind patterns that
drive sea ice off-shore (Schofield et al., 2010). The reduced
sea-ice duration lengthens the growing season for water-
column phytoplankton, which are shaded from sunlight
by sea-ice cover for much of the year. Further, sea-ice
melting during the spring releases freshwater, acting to
cap the water column, shallow the ocean mixed layer depth,
and promote surface plankton blooms by trapping plankton
in a well-lit surface layer. Time series satellite ocean color
records for the northern regions of the west Antarctic
Peninsula indicate that phytoplankton stocks have declined
by over 80% because there is less sea ice to stabilize the
upper water column and because of increased wind-driven
turbulence (Montes-Hugo et al., 2009). In contrast, phyto-
plankton blooms have increased in the south in the
present-day seasonal marginal ice zone because more light
is penetrating the ocean as the sea ice declines.
The decline of sea ice and warmer, more maritime
weather conditions may be the cause of the major eco-
system regime shifts now being observed around the Ant-
arctic Peninsula (Figure 31.1). Impacts can be especially
strong for organisms adapted to polar conditions that
require the presence of sea ice for part of their life history.
Ice-dependent species include krill, relatively large plank-
tonic crustaceans that are quite abundant in Antarctic waters
and a critical prey for seabirds and marine mammals. Small
juvenile krill over-winter under the ice pack, hiding from
predators and grazing on ice algae. Krill have declined by
an order of magnitude in the Atlantic sector of the Southern
Ocean since 1950 with a corresponding increase in the
abundance of salps, a gelatinous tunicate filter feeder
(Atkinson et al., 2004). Salps are not a very palatable food
source for most marine predators, and therefore, a switch
from krill to salps could lead to important consequences
for higher trophic levels.
Sea ice also provides an important habitat for many
polar seabirds and mammals (e.g., penguins, polar bears,
walruses, seals) that use the ice as a foraging platform or
breeding habitat, suggesting that these species will face
problems with warming. Under warming conditions, the
spatial range of the polar species will likely contract toward
the pole, whereas for subpolar species adapted to warmer
conditions, this may allow migration into newly available
habitats. Temporal mismatch between prey availability
and predator food demands, which are often constrained
1970
15,000
10,000
5000
01980 1990
Year
Ade
lie b
reed
ing
pairs
2000 2010
3000
Gentoo and C
hinstrap breeding pairs
2000
1000
0
FIGURE 31.7 Time series of penguin populations in the
local area around Palmer Station, Antarctica. The popu-
lation is estimated from the number of breeding pairs mea-
sured during the summer breeding season. Three species
are shown, polar ice-dependent Adelie (red; left-hand axis)
and subpolar Chinstrap (blue) and Gentoo (black) (both on
right-hand axis). Over time, in response to sea-ice decline
and increased snow, the polar Adelie population has
declined by about 80%, replaced by an influx from the
north of Chinstrap and Gentoo penguins that are ice intol-
erant. Data and images courtesy of William Fraser and
Hugh Ducklow.
PART VI The Changing Ocean826
by aspects of their life history, can be especially prob-
lematic for polar and migratory seabirds that depend upon
an ample and reliable food supply of krill and small fish
during the limited summer breeding season. Along the Ant-
arctic Peninsula, the population of Adelie penguins has
declined by 80% in the Palmer Station region because of
cascading responses to sea-ice loss, reduced food avail-
ability, and elevated late-spring snowfalls (Figure 31.7;
Schofield et al., 2010). Similar regional declines have been
observed in crab-eater seals, another ice-dependent species.
Conversely, in a case of opening niche space and species
expansion, ice-intolerant gentoo and chinstrap penguins,
as well as southern fur seals, perhaps, are now migrating
into the region and establishing new breeding colonies. In
another example of a climate-related invasive event with
broad ecological impacts, warming of bottom waters along
the continental slope and shelf appears to be allowing the
recent and ongoing colonization of the Antarctic Peninsula
by king crab, shell-crushing predators that have been absent
from the Antarctic food web for about 25 million years
(Fox, 2012; Smith et al., 2012).
In the Arctic and adjacent marginal seas, emerging bio-
logical–climate signals include: marine species range shifts;
changes in abundance, growth, condition, behavior, and phe-
nology of some species; and community and regime shifts
(Wassmann et al., 2011). An unusually large fraction of
primary production in the seasonally ice-covered northern
Bering shelf ecosystem sinks to the seafloor and supports a
large and diverse benthic community; pelagic fish predation
is limited by cold-water temperatures and ice cover, allowing
diving seabirds, bearded seals, walrus, and gray whales to
harvest the high benthic production (Grebmeier et al.,
2006). Warming and variability in sea-ice retreat coincide
with declines in clam populations, which in turn co-occur
with dramatic declines in diving sea ducks, more northerly
migrations of large vertebrate predators (walrus and gray
whales), and potentially poleward-expanding ranges for
pelagic fishes (Grebmeier et al., 2010).
In temperate oceans, poleward range shifts are evident
for many fish species based on long time series from com-
mercial fish stock surveys (Perry et al., 2005). Nye et al.
(2009), for example, analyzed temporal records of fish
species’ distributions in the Northeast United States conti-
nental shelf ecosystem. The southern stocks tended to move
northward in time in response to warming, appearing to
maintain an approximately constant thermal habitat
(Figure 31.8). In the more geographically restricted Gulf
of Maine, northern stock data indicated no significant lati-
tudinal migration but rather movement toward deeper and
colder waters.
The concept that species’ ranges are bounded within a
fixed range of environmental properties is the basis for bio-
climate envelope models that can be used to project fish
species responses to climate change including local
extinction and species invasion (Cheung et al., 2009).
Habitat suitability, defined using present-day biogeo-
graphic distributions and environmental conditions, is
folded into dynamic population models that capture growth,
larval dispersal, and adult migration. Cheung et al. (2009)
forecast substantial reorganization of global fishery catch
potential with large declines in the tropics and increases
in high-latitude regions, overlain with large regional varia-
tions. Refinements to this class of models and other
methods are being used to examine likely future changes
in fish body size with warming, in interactions with ocean
acidification and declining O2, in interactions of climate
effects with overfishing, and in impacts on the economics
of fisheries (Brander, 2007, 2010; Sumaila et al., 2011).
However, shifts in species’ geographic ranges depend
on much more than simply temperature, and good fore-
casting likely will require an understanding of the life
history of particular species. For example, while warming
may open up new habitat in polar Arctic regions for cod,
herring, and pollock (Loeng et al., 2005), the continued
presence of cold bottom-water temperatures on the shelf
could limit northward migration into the northern Bering
–8
1.6
0.07
0.03
0
–0.03
0.08
0.04
–0.04
–0.08–0.07
0.8
0
–0.4
–1.6
–0.2 0 0.2
–0.5 0 0.5
–0.2 0 0.2
Proportion of stocks
Slo
pe o
f lin
ear
regr
essi
on
–0.4 0 0.4
–6
0
6
810
6
0
–6
–10
0.06
0.04
–0.04
–0.06–0.4 0 0.4
–0.4 0 0.4
p = 0.007
p = 0.004
(a) Poleward shift (distance) (d) Area occupied
(e) Maximum latitude(b) Mean depth
(c) Mean temperature (f) Minimum latitude
FIGURE 31.8 Histograms comparing distributional responses for southern (red) and northern (blue) stocks in (a) distance shifted poleward, (b) mean
depth, (c) mean temperature, (d) area occupied, (e) maximum latitude, and (f) minimum latitude. Significant differences detected with a Mann–Whitney
U-test between species found in the two ecoregions are indicated by p values. From Nye et al. (2009).
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 827
Sea and Chukchi Sea (Sigler et al., 2011). In addition,
warming may cause reductions in the abundances of some
species, such as pollock, over their current ranges in the
Bering Sea (Mueter et al., 2011), linked to changes in
overall food-web dynamics and bottom-up food resources,
not just direct thermal effects (Hunt et al., 2011). Climate
change information needs to be incorporated into man-
agement strategies for specific fisheries (Ianelli et al.,
2011) and assessments of the vulnerabilities of keystone,
sentinel, iconic, and endangered marine species to local
population collapse. Wolf et al. (2010) illustrate such an
approach integrating climate trends into a demographic
model for Cassin’s auklet, a seabird that feeds on plankton
and a sentinel species in the California Current System.
More broadly, climate model projections are increasingly
being used in a variety of ways to evaluate climate impacts
on living marine resources (Stock et al., 2011).
Climate variations and climate change also influence the
spread and impact of marine diseases and parasites (Harvell
et al., 2002). Marine disease appears to be on the rise with
time, and higher SSTs have been linkedwith higher intensity
and increased spatial ranges of diseases that attack corals,
abalones, oysters, fishes, and marine mammals (Ward and
Lafferty, 2004). Climate warming acts through several dif-
ferent pathogen-specific mechanisms. Warming can
increase pathogen over-wintering survival, tied to the
northward spread of Dermo disease, an oyster parasite, up
theU.S. east coast (Cook et al., 1998) and the growth of coral
disease lesions (Weil et al., 2009). Higher seasonal temper-
atures may cause an expansion ofVibrio species, pathogenicbacteria that infect oysters, and may cause human illness
(Baker-Austin et al., 2013). Warming can also increase
pathogen susceptibility by intensifying thermal stress
because of the elevated size and duration of positive temper-
ature anomalies. Record warm tropical SSTs have caused
widespread coral disease outbreaks (Miller et al., 2009)
and coral bleaching (Eakin et al., 2010). Bleaching occurs
in response to environmental stress when the naturally col-
orless coral polyps expel their zooxanthellae, the colored
symbiotic dinoflagellates whose photosynthesis fuels the
growth of their coral hosts (Figure 31.9).
In fact, coral reefs are some of the most susceptible eco-
systems to climate warming because of the sensitivity of
coral-algal symbiosis to minor increases in maximum sea-
sonal temperature; warming of as little as 1 �C can cause
coral bleaching (Hoegh-Guldberg et al., 2007; Donner,
FIGURE 31.9 A colony of star coral (Montastraea faveolata) off the southwest coast of Puerto Rico, estimated to be about 500 years old, exemplifies the
effect of rising water temperatures. Increasing diseases due to warming waters (a) were followed by such high temperatures that bleaching or loss of
symbiotic microalgae from coral occurred (b), followed by more disease (c) that finally killed the colony (d). Figure courtesy of Ernesto Weil.
PART VI The Changing Ocean828
2009). Warm-tolerant zooxanthellae may become more
predominant on reefs in the future, allowing some corals
to survive moderate temperature increases, but this may
have negative impacts on other aspects of coral health such
as growth (Jones and Berkelmans 2010). Reefs are also
threatened by a variety of local human pressures including
pollution, overfishing, and dredging, and the degradation of
reefs affects not only corals themselves but also the rich bio-
diversity of other organisms living within the structural
complexity of the coral reef seascape. Coral reefs are
important for human societies, often supporting locally
essential artisanal reef fisheries, and reefs are some of the
most valuable marine ecosystems because of tourism and
recreation income and coastal protection (Cooley et al.,
2009). Ocean acidification due to rising atmospheric CO2
is an additional threat to corals and other important reef cal-
cifying organisms such as crustose coralline algae that build
reef frameworks (Anthony et al., 2008).
4. OCEAN ACIDIFICATION
Climate change is not the only important effect of rising
atmospheric CO2, which also causes direct changes in sea-
water acid–base and inorganic carbon chemistry, termed
ocean acidification that can impact marine organisms and
ecosystems. Compared to most freshwater systems, the
acid–base chemistry of seawater is relatively stable because
the inorganic carbon system and large alkalinity levels
buffer seawater pH. Many marine organisms appear to be
adapted to relatively constant local acid–base conditions
and are sensitive to relatively small variations in pH and
the concentrations of various inorganic carbon species
(Doney et al., 2009; Gattuso and Hansson, 2011). The ocean
uptake of anthropogenic CO2 is causing global-scale shifts
in upper-ocean chemistry that are rapid compared to varia-
tions in the geological past; for example, the surface ocean
pH change caused by the �30% rise in atmospheric CO2
associated with the last deglaciation was roughly two orders
of magnitude slower than the current rate driven largely by
fossil-fuel burning (Honisch et al., 2012). The chemistry of
ocean acidification is relatively well understood (Feely
et al., 2009); biological implications are slowly becoming
clearer at the level of individual species, but substantial
uncertainties remain particularly at the ecosystem level
(Gattuso et al., 2011).
CO2 acts as a weak acid when added to water at seawater
pH levels:
CO2þH2O,HþþHCO�3 ð31:4Þ
The forward reaction releases hydrogen (Hþ) and bicar-
bonate (HCO�3 ) ions and lowers pH, defined as
pH¼�log10{Hþ}. Most of the extra Hþ ions react with
and lower CO2�3 concentrations:
HþþCO2�3 ,HCO�
3 ð31:5ÞCO2 input also increases aqueous CO2(aq) and DIC, [DIC]¼
[CO2(aq)]þ [HCO�3 ]þ [CO2�3 ]. Another important reaction
is the dissolution or precipitation of solid CaCO3 used by
many marine plants and animals to form shells and hard
body parts:
CaCO3 sð Þ,Ca2þþCO2�3 ð31:6Þ
CaCO3becomesmore soluble asCO2 rises andCO2�3 declines,
representedmathematically by a loweringof theCaCO3(s) sat-
uration state,O¼ [Ca2þ] [CO2�3 ]/Ksp,whereKsp is the thermo-
dynamic solubility product that varies with temperature,
pressure, and mineral form. Present-day ocean surface waters
are currently supersaturated (O>1) for the two major forms
used by marine organisms, the more soluble form, aragonite
(corals,manymollusks), and the lesssoluble form,calcite (coc-
colithophores, foraminifera, and some mollusks). Other rela-
tively soluble forms include amorphous CaCO3 and CaCO3
containing various amounts of magnesium.
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 829
Natural physical and biological processes influence sea-
water acid–base chemistry, leading to large-scale spatial
gradients and seasonal variability in pH, O, and inorganic
carbon speciation (Feely et al., 2009). High-frequency tem-
poral variability on diurnal to weekly timescales is also
observed in coastal, estuarine, and coral reef systems
(Hofmann et al., 2011). In general, surface waters tend to
have lower CO2 and DIC and, therefore, slightly higher
pH because of phytoplankton uptake of inorganic carbon
as part of photosynthesis. The opposite pattern and low
O2 values are found in the thermocline because of the res-
piration of sinking organic matter and downward trans-
ported dissolved organic matter. CaCO3 saturation state
decreases with depth because of organic matter respiration
and pressure effects on CaCO3 solubility. Aragonite and
calcite often become undersaturated (O<1) below some
depth in the water column, at which point, unprotected
shells and skeletons begin to dissolve; the saturation depth
horizon is particularly shallow in the Pacific that has a high
burden of metabolic CO2 (Feely et al., 2009). Because of
increased CO2 solubility and temperature effects on the
thermodynamic equations, cold polar surface waters exhibit
lower CO2�3 ion concentrations and O values.
Long-term trends in pH and inorganic carbon chemistry
are clearly evident over the past several decades in ocean
time series and hydrographic surveys (Dore et al., 2009;
Byrne et al., 2010; Bates et al., 2012). From preindustrial
levels, contemporary surface ocean pH is estimated to have
dropped on average by about 0.1 pH units (a 26% increase
in [Hþ]), and further decreases of 0.2 and 0.3 pH units will
occur over this century unless anthropogenic CO2 emis-
sions are curtailed dramatically (Orr et al., 2005). Surface
FIGURE 31.10 Summary plot of variations in calcification rate, normalized
corals. Figure courtesy of Chris Langdon.
ocean CaCO3 saturation state is declining everywhere,
and model simulations indicate that polar surface waters
will become undersaturated for aragonite when atmospheric
CO2 reaches 400–450 ppm for the Arctic and 550–600 ppm
for the Antarctic (Orr et al., 2005; Steinacher et al., 2009).
Because of the larger natural background CO2 levels, sub-
surface waters have a lower buffer capacity and exhibit a
larger pH drop per amount of CO2 added; this increases
the susceptibility to acidification of O2 minimum zones
(Brewer and Peltzer, 2009), coastal waters that are already
experiencing nutrient eutrophication and hypoxia (Feely
et al., 2010; Cai et al., 2011), and eastern boundary current
upwelling systems (Feely et al., 2008; Gruber et al., 2012).
Numerous biological effects have been measured in
response to ocean acidification for both pelagic
(Riebesell and Tortell, 2011) and benthic (Andersson
et al., 2011) organisms (see also Fabry et al., 2008;
Doney et al., 2009). Most biological impacts have been
inferred from short-term manipulation experiments at the
organism level to step-increases in CO2, for example, lower
calcification rates in corals and mollusks, higher photosyn-
thesis rates for seagrasses and some phytoplankton groups,
and increased nitrogen fixation by cyanobacteria. For
example, Figure 31.10 shows a summary plot for tropical
corals of the relative decrease in calcification rate as a
function of declining CaCO3 saturation state.
Organism responses may vary with life history stage,
with juveniles often more susceptible than adults, and some
organisms may be able to accommodate elevated CO2 but at
an additional energetic cost with consequences for devel-
opment, reproduction, and fitness. Different organism groups
may be sensitive to different aspects of seawater chemical
to preindustrial values in percent, to aragonite saturation state for tropical
-0.8
Surviv
al
Calcific
ation
Growth
Photo
synt
hesis
Develo
pmen
t
Abund
ance
Met
aboli
sm
-0.6Mea
n ef
fect
siz
e (L
nRR
)
-0.4
-0.2
0
0.2
(82)
(173)
(110)
(69)
(24) (72)
(32)
∗∗
∗
∗
∗
FIGURE 31.11 Summary results for a meta-analysis of biological
impacts of ocean acidification reported in the literature. Key physiological
responses are aggregated across all taxa. The effect size is the ratio of the
mean effect in the acidification treatment to the mean effect in a control
group and is scaled to �0.5 unit reduction in pH. Error bars represent
95% confidence intervals. Asterisks denote statistically significant effects,
and the number of studies is shown in parentheses. Figure from Kroeker
et al. (2013).
PART VI The Changing Ocean830
trends: for calcifiers declining CO32� levels; for autotrophs,
increased aqueous CO2; and for adult fish and cephalopods,
acid–base regulation and CO2/O2 transport and gas
exchange. In a recent meta-analysis of available literature
studies, Kroeker et al. (2010, 2013) show that statistically
significant declines are observed for survival, calcification,
growth, development, and abundance, with substantial vari-
ations across taxonomic groups (Figure 31.11).
Effects on natural populations and communities so far
have been more difficult to detect outside of a limited number
of pelagicmesocosm experiments and some studies in isolated
high-CO2 environments such as shallow volcanic vents that
tend to support laboratory findings (Hall-Spencer et al.,
2008; Fabricius et al., 2011). Shellfish hatcheries along the
Oregon–Washington coast have experienced dramatic
declines in oyster harvest (Barton et al., 2012) in response
to the upwelling onto the shelf of strongly acidified waters
with lowCaCO3 saturation state (Feely et al., 2008). In general
though, the nature and magnitude of the responses within
natural populations and the ability of organisms to acclimate
or adapt to gradual CO2 trends are still mostly unknown.
5. DEOXYGENATION AND HYPOXIA
Marine biota can also be influenced by ocean oxygen dis-
tributions, particularly at low O2 values in oxygen
minimum zones and hypoxic coastal environments (Levin
et al., 2009; Keeling et al., 2010). Dissolved O2 gas is
required for aerobic respiration, and below certain
organism-specific thresholds, low O2 begins to affect
metabolic rates and behavior. Low O2 leads to marine
habitat degradation and, in extreme cases, extensive fish
and invertebrate mortality, and larger mobile animals often
move out of low oxygen environments, resulting in so-
called dead-zones where many macrofauna are nearly
absent (Diaz and Rosenberg, 2008; Rabalais et al., 2010).
Thresholds for hypoxia vary by organism but are typically
�60 mmol kg�1 or about 30% of surface saturation. Under
suboxic conditions (<5 mmol kg�1), microbes begin to
utilize nitrate (NO3�) rather than O2 as the terminal electron
acceptor for organic matter respiration (leading to denitrifi-
cation), resulting in reactive nitrogen loss and N2O pro-
duction. By simultaneously removing O2 and adding
CO2, organic matter respiration can induce multiple
stressors on organism physiology, and O2 stress can also
reduce organism thermal tolerances (Portner and Farrell,
2008; Portner et al., 2011).
Hypoxic conditions occur naturally in open-ocean and
coastal subsurface waters from a combination of weak ven-
tilation, warming, and organic matter degradation—
features that are all exacerbated by climate warming and
coastal nutrient eutrophication. Coastal hypoxic systems
are widespread globally with more than 400 instances cov-
ering an area >245,000 km2 (Diaz and Rosenberg, 2008).
An expansion in the duration, intensity, and extent of
coastal hypoxia over the past several decades is attributed
to growing coastal urbanization, land runoff of excess
nutrients from fertilizers and sewage, and atmospheric
nitrogen deposition from fossil-fuel combustion. About half
the global riverine nitrogen input is anthropogenic in origin
(Seitzinger et al., 2010), and coastal nutrient eutrophication
is also associated with increased frequency of harmful algal
blooms (Anderson et al., 2002). The emergence of hypoxia
in other coastal regions may be related to variations or
trends in ocean-atmospheric physics. Increased wind-
driven upwelling is linked to the first appearance of hypoxia
and even anoxia on the inner-shelf off Oregon–Washington
coast after five decades of hypoxia-free observations (Chan
et al., 2008).
Oxygen minimum zones occur naturally in the open-
ocean in the tropics and subtropics, with the lowest O2
suboxic waters restricted to the Arabian Sea, the Bay of
Bengal, and the eastern tropical Pacific (Paulmier and
Ruiz-Pino, 2009; Bianchi et al., 2012). In most subtropical
thermocline regions, oxygen levels tend to be relatively
high because the anticyclonic gyre circulation transports
oxygen-rich water relatively rapidly and directly from
surface isopycnal outcrops into the interior. In the tropics,
low oxygen regions typically occur equatorward of these
directly ventilated subtropical waters in regions more cut
off from the atmosphere where ventilation occurs more
indirectly through diffusive processes, so-called shadow
zones of the wind-driven circulation. Oxygen minimum
zones often underlie biologically productive regions such
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 831
as the eastern boundary current upwelling systems (e.g.,
Peru, Benguela). The lowest water-column O2 values are
observed in the upper and mid-thermocline where biological
oxygen consumption rates are high. A number of studies
indicate that subsurface oxygen values are declining with
time in the midlatitudes (Whitney et al., 2007; Helm et al.,
2011) and that oxygen minimum zones are expanding
spatially in both the vertical and horizontal directions
(Stramma et al., 2008), which could limit the habitat range
for some fish species and alter diel vertical migration patterns
(Stramma et al., 2012). Only a fraction of the oxygen loss is
directly related to warming and decreased O2 solubility, and,
therefore, most of the deoxygenation must reflect alterations
and slowdowns of thermocline ventilation. In response to
climate warming over the twenty-first century, model projec-
tions indicate further reductions in the global oxygen
inventory and expansions of open-ocean oxygen minimum
zones (Bopp et al., 2002; Frolicher et al., 2009).
6. MARINE BIOGEOCHEMICALCYCLES–CLIMATE INTERACTIONS
Marine biota and especially microbes can affect the compo-
sition of the atmosphere by producing and consuming a
number of trace gases that influence climate and atmospheric
chemistry (Denman et al., 2007). Ocean sources and sinks of
radiatively active trace gases such as CO2, N2O, methane,
and dimethylsulfide (DMS) are of particular interest because
they can mediate biological–climate feedbacks and the
amplification or damping of external climate perturbations
(Boyd and Doney, 2003; Doney and Schimel, 2007).
Methane, CO2, and N2O are powerful “greenhouse” or
heat-trapping gases in the atmosphere, absorbing far-
infrared or longwave radiation emitted by the land and
ocean surface; an increase in atmospheric greenhouse gases
thus leads to surface heating and climate warming (Kiehl
and Trenberth, 1997; Solomon et al., 2007). Positive feed-
backs with other components of the physical climate system
(e.g., sea ice and snow cover) tend to amplify the heating
caused by the initial biogeochemical radiative perturba-
tions, leading to further warming; key ocean climate feed-
backs include warming SSTs that increase evaporation and
elevate atmospheric water vapor, another powerful green-
house gas, and retreating bright, high-albedo sea ice that
exposes dark, low-albedo waters that absorb more solar
radiation.
Atmospheric CO2 and climate have coevolved with the
biosphere over the history of the Earth, with periods of ele-
vated CO2 generally reflected in warmer global climate
conditions (Siegenthaler et al., 2005; Doney and Schimel,
2007; Royer et al., 2012). The net balance of sources and
sinks that determines atmospheric CO2 is sensitive to
climate variations, and climate warming may both result
from and cause higher atmospheric CO2 levels. Ocean cir-
culation and biogeochemistry are major factors governing
atmospheric CO2 (Denman et al., 2007; see also
Chapter 30 on the carbon cycle; Tanhua et al. 2013). The
large ocean inventory of inorganic carbon (�37,100 PgC)
is roughly 50 times the CO2 inventory in the preindustrial
atmosphere (�590 PgC), and the ocean carbon pool is
the largest mobile reservoir on the planet on timescales
of decades to millennia (Sarmiento and Gruber, 2002).
Volk and Hoffert (1985) described a simple conceptual
model for how solubility and biological processes or
“pumps” affect the vertical redistribution of inorganic
carbon within the ocean and thus total ocean carbon storage.
CO2 solubility is temperature dependent; warm water holds
less DIC, and climate warming will therefore reduce ocean
carbon storage. The biological pump consists of two com-
ponents, the sinking fluxes of organic matter and inorganic
CaCO3 that transport carbon from the surface ocean to
depth. Organic carbon production lowers surface water
CO2, driving a net CO2 uptake from the atmosphere; in con-
trast, CaCO3 shell formation reduces surface water alka-
linity more than it reduces DIC and, thus, effectively
causes a net efflux of CO2 to the atmosphere.
As recorded in Antarctic ice cores, atmospheric CO2
levels underwent large variations over glacial–interglacial
cycles with low values of �180 ppm during cold glacial
maxima and high values of �280 ppm during warm inter-
glacial periods (EPICA community members, 2004). The
large atmospheric CO2 variations must reflect reorganiza-
tions of ocean circulation and biogeochemistry, leading to
substantial changes in ocean carbon storage (Sigman and
Boyle, 2000). During the last deglaciation, paleo-evidence
indicates that initial northern hemisphere warming led to a
reduction in Atlantic meridional overturning circula-
tion that, in turn, triggered Southern Ocean CO2 release
and global warming (Shakun et al., 2012a,b). Numerous
hypotheses have been proposed to explain glacial–
interglacial ocean CO2 variations associated with circu-
lation, biological productivity, and sea-ice effects;
recent work argues that CO2 degassing was related to
increases in Southern Ocean upwelling (Anderson et al.,
2009). Atmospheric CO2 variations have been considerably
smaller during the recent warm Holocene (the past
�11,000 years) with a rise of only �20 ppm over the past
7000 years attributed to a mix of terrestrial and oceanic
processes including shallow-water carbonate deposition
(coral reefs) and slow adjustment of deepwater chemistry
and sediments (Menviel and Joos, 2012) (Figure 31.12).
Since the preindustrial era (i.e., since�1800 CE), atmo-
spheric CO2 levels have increased by more than 40% from a
preindustrial level of approximately 280 ppm to 395 ppm at
Mauna Loa observatory by the end of 2012 (Tans and
Keeling, 2012). Based on isotopic composition and detailed
carbon budgets, this CO2 rise can be tied definitively to
800-450
-420
-390
180
210
240
270
300
600 400
Age (kyr)
Atmospheric CO2 over late Pleistocene
∂D ic
e (‰
)C
O2
(ppm
v)
200 0
and Dome C dataVostok
EPICA Dome C data
FIGURE 31.12 Glacial–interglacial variations of atmospheric CO2 (ppmv) and ice deuterium (dD in %), a proxy for temperature (higher dD reflects
warmer conditions), from Antarctic ice cores for the past 650,000 years. Figure from Doney and Schimel (2007).
PART VI The Changing Ocean832
human activities, in particular deforestation, fossil-fuel
combustion, and cement manufacture. The ocean has
played a critical climate service by removing some of the
excess or anthropogenic CO2 from the atmosphere
(Sabine et al., 2004; Sabine and Tanhua, 2010). Estimated
ocean carbon uptake in 2008 was 2.3�0.4 PgC year�1
compared to a fossil-fuel combustion release to the atmo-
sphere of 8.7�0.5 PgC year�1 (Le Quere et al., 2009).
Cumulative ocean carbon uptake since the beginning of
the industrial age is equivalent to about 25–30% of total
human CO2 emissions (Sabine and Tanhua, 2010). The
global ocean uptake rate is governed primarily by the atmo-
spheric CO2 excess and trend and the rate of ocean circu-
lation that exchanges surface waters equilibrated with
elevated CO2 levels with subsurface waters that have not
yet been exposed to the anthropogenic CO2 transient
(Sarmiento et al., 1992; Khatiwala et al., 2009).
Ocean carbon storage is enhanced when more of sub-
surface nutrient inventory is biologically released rather
than “preformed,” the latter component referring to
nutrients that are advected into the ocean interior from
nutrient-rich surface waters (Ito and Follows, 2005). The
largest reservoir of unused surface macronutrients resides
in the Southern Ocean, and modeling studies suggest that
ocean carbon storage is especially sensitive to Southern
Ocean deepwater formation (Marinov et al., 2006). A
number of other biogeochemical factors can enhance ocean
carbon storage, and many of these mechanisms are sensitive
to physical climate (Boyd and Doney, 2003; Denman et al.,
2007). Warmer stratified conditions or increased iron inputs
via dust deposition could promote subtropical nitrogen fix-
ation, increasing the pool of bioavailable nitrogen (Boyd
and Doney, 2002; Moore et al., 2006). Atmospheric iron
deposition could also increase the extent of surface nutrient
utilization in HNLC areas (Martin, 1990). Ocean acidifi-
cation and climate-driven shifts in community composition
may increase the carbon to nutrient and organic carbon to
CaCO3 stoichiometry of sinking export material (Oschlies
et al., 2008; Gehlen et al., 2011). Plankton community shifts
could also increase (or perhaps decrease) the vertical length
scale over which sinking organic matter is regenerated
(Kwon et al., 2009).
Climate warming is projected to reduce ocean uptake of
anthropogenic CO2 due to decreased solubility, increased
vertical stratification, and slowing of intermediate and deep-
water formation in the North Atlantic (Sarmiento and Le
Quere, 1996; Friedlingstein et al., 2006). The ocean also
becomes less efficient with time at removing further atmo-
spheric CO2 because ocean acidification lowers seawater
buffer capacity. Climate-governed changes in ocean vertical
exchange also slow the nutrient supply to the surface
resulting in reduced biological carbon export. In principle
this should lower ocean CO2 uptake even further; however
this effect is counteracted by the reduced upward flux ofmet-
abolic CO2, and the net climate effect on the biological pump
is a modest increase in the effective carbon sink (Fung et al.,
2005). Strengthening of the westerly winds in the Southern
Ocean may be increasing vertical upwelling of CO2-rich Cir-
cumpolar Deep Water, which would increase the ocean
efflux of natural CO2, reducing the global net anthropogenic
CO2 uptake (Le Quere et al., 2007; Lovenduski et al., 2008).
There is some evidence that anthropogenic climate change is
already slowing ocean CO2 uptake (Le Quere et al., 2010).
However, detecting a climate signal trend is difficult given
the large seasonal and spatial variations in ocean CO2
uptake and release (Takahashi et al., 2009) and the sub-
stantial interannual to interdecadal variability in air–sea
CO2 flux driven by natural climate modes, in particular
ENSO (Park et al., 2010).
Ocean acidification may alter the ocean carbon cycle via
impacts on the export flux and subsurface remineralization
for either CaCO3 or organic matter. The net effect on ocean
Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 833
carbon storage varies with both positive and negative feed-
backs but is relatively small in current models. Increased
carbon to nutrient ratios in sinking organic matter seen in
some mesocosm experiments exposed to high CO2 could
expand subsurface low oxygen zones and increase N2O pro-
duction (Oschlies et al., 2008).
N2O is produced by marine microbes as a minor by-
product of two biogeochemical pathways involved with
organic matter respiration, nitrification (NHþ4 !NO�
3 )
throughout most of the water column, and denitrification
(NO�3 !N2) restricted to hypoxic and suboxic waters
(Codispoti, 2010). The traditional paradigm suggested that
the two processes contribute about equal amounts to global
N2O production, though there is growing evidence that
nitrification dominates (Dore et al., 1998; Freing et al.,
2012). Subsurface ocean N2O distributions are correlated
with apparent oxygen utilization (AOU¼ [O2]saturation�[O2]measured), consistent with observations that the nitrifi-
cation yield of N2O per mole NO�3 produced increases with
declining O2 (Nevison et al., 2003). Under climate warming
scenarios, areal expansion of suboxic waters will likely
increase marine denitrification but may not substantially
alter ocean N2O production (Bianchi et al., 2012), particu-
larly if denitrification is a minor global source. On the other
hand, broad-scale deoxygenation in the ocean thermocline
could increase N2O yield from nitrification. This may be
partially countered by ocean acidification, which has been
shown to slow microbial nitrification (Beman et al., 2011).
The net effect is as yet uncertain.
Surface ocean DMS levels and air–sea fluxes are gov-
erned by a complex set of food-web interactions: phyto-
plankton production of nongaseous organic sulfur
precursors; biological cleavage to DMS; bacterial and pho-
tochemical DMS destruction (Toole et al., 2008). In tem-
perate waters, elevated DMS is associated with increased
primary production and particular organosulfur-rich phyto-
plankton species; in subtropical and tropical regions, high
DMS production is associated with ultraviolet and low-
nutrient stress, both of which could result from warming
and increased vertical stratification (Toole and Siegel,
2004). Modeling studies suggest climate warming may
increase Southern Ocean DMS fluxes to the atmosphere
because of sea-ice retreat and changes in phytoplankton
composition (Cameron-Smith et al., 2011) though current
model DMS parameterizations may be insufficient to make
robust climate change projections (Halloran et al., 2010).
When released to the atmosphere, biologically produced
DMS can form aerosols and cloud condensation nuclei in
the remote marine atmosphere. In a seminal paper,
Charlson et al. (1987) argued for a biological-climate reg-
ulation mechanism, by which warming would increase
DMS flux to the atmosphere, leading to more marine stratus
cloud cover and surface cooling. The so-called CLAW
hypothesis, named after the authors of Charlson et al.,
involves a complex suite of biological and chemical steps
(Vogt and Liss, 2009), and some researchers argue that,
after more than two decades, there is little evidence to
support biological control over cloud condensation nuclei
levels and that the CLAW hypothesis should be abandoned
(Quinn and Bates, 2011).
7. OBSERVATIONAL AND RESEARCHDIRECTIONS
Studying climate-ecosystem dynamics in the ocean is quite
challenging because of the long timescales and large space
scales involved as well as the complexity of marine food
webs that span from viruses and bacteria to apex predators.
In many cases, we know considerably more about the direct
responses of a particular species to short-term physical var-
iations and have to infer longer time-scale effects. Or we
may have time series data for ecosystem parameters or
species abundance but do not fully understand the under-
lying mechanisms or the indirect feedbacks on other species
through altered food webs. Individual research techniques
each have their own strengths and weaknesses, and a mix
of different, complementary approaches is required com-
bining observations, experiments, theory, and modeling.
The time scales required to resolve climate–ecosystem
interactions are inherently multiannual to multidecadal,
and most available ocean field datasets are of insufficient
duration. The situation is improving with time, but estab-
lishing and maintaining long-term observational records
should be a top priority for the ocean research community.
In many cases, observing the ecological response to inter-
annual and decadal variability on a regional scale may
inform estimates of future climate-driven biological trends
that are more difficult or impossible to address through
other approaches (Boyd and Doney, 2002).
Many key insights have been derived by co-opting time-
series data that were originated for other purposes such as
surveying commercial fishery stocks or addressing specific
process-oriented questions. Problems may arise related to
standardization of measurements, data quality, and data
continuity when data were not collected with climate-scale
analysis in mind. Even with such caveats, the availability of
long-term ecological data is invaluable. Commonly used
records include local and regional datasets such as
CALCOFI, Joint Global Ocean Flux Study time series,
the Atlantic Meridional Transect line, Long-Term Eco-
logical Research (LTER) sites, and the CPR (Ducklow
et al., 2009). For the upper-ocean, satellite data records of
ocean color only began with the CZCS in the late 1970s,
and the Sea-viewingWide Field-of-view Sensor (SeaWiFS)
provided more than a decade of nearly continuous global
coverage of surface chlorophyll and primary productivity
from late 1997 through 2010 (McClain, 2009; Siegel
PART VI The Changing Ocean834
et al., 2013). Efforts are also underway to deploy more
extensive global in situ observing systems that will inform
marine ecology and biogeochemistry. Major advances have
been made on integrating bio-optical and chemical sensors
on new and autonomous observing platforms including:
moorings, profiling floats, drifters, subsurface gliders, wave
gliders, and AUVs (Johnson et al., 2009), and observational
plans have been developed, for example, to add O2, pH, and
bio-optical sensors to the Argo float array (Gruber et al.,
2010) or to track ocean acidification and subsequent bio-
logical impacts in open-ocean and coastal regions
(Iglesias-Rodriguez et al., 2010).
Numerous process studies have been conducted on lab-
oratory cultures and field samples to access the sensitivity
of organisms and biological communities to temperature,
CO2, pH, O2, nutrients, trace metals, and other environ-
mental factors (Somero, 2012). However, some care must
be taken in the interpretation of these results and their
extrapolation to longer-term and ecosystem-level impacts,
especially for short-duration perturbation experiments
that simulate climate change from the acute biological
responses to large, abrupt environmental changes.
Process-based studies will remain an essential tool for pro-
viding a mechanistic framework to explain the phenomeno-
logical signals and trends in field observations, and moving
forward, more emphasis is needed on the synergistic effect
of multiple stressors (warming, lower pH, etc.) that will
occur contemporaneously in the future (Boyd et al.,
2008). Also, more emphasis is needed on the ecological
resilience of marine systems (Bernhardt and Leslie, 2013)
and microevolution in response to climate change, espe-
cially for planktonic organisms with relatively short gener-
ation time scales (Dam, 2013). Considerable insights have
been derived from mesocosm experiments on planktonic
communities in which various environmental parameters
are manipulated either individually or in a factorial fashion.
Open-ocean iron fertilization experiments have been a great
success (Boyd et al., 2007), and emerging technologies,
such as wave pumps (White et al., 2010) and free-ocean
CO2 release may allow for other types of ocean manipu-
lation experiments (Kline et al., 2012).
A deeper mechanistic understanding is also desirable for
deciphering the past and predicting the future using
numerical models. Although statistical relationships can
be deduced relating biological response to climate forcing,
these relationships may not hold outside the bounds of
present-day conditions. Over the twenty-first century,
anthropogenic climate change may be substantial enough
to create no-analogue or novel ocean ecosystem states,
where, because of spatial range shifts and changes in abun-
dance, the biological community does not match well any
present-day system. Prognostic modeling can be a powerful
tool for projecting into an uncertain future but only if the
model dynamics is adequately known and the model tested
thoroughly against available data (Glover et al., 2011).
Marine ecology and biogeochemistry are increasingly being
incorporated into basin and global ocean general circulation
models as well as coupled ocean-atmosphere climate
models, and substantial progress has been made from only
a decade ago (Doney, 1999; see also Chapter 26 on
modeling ocean biogeochemistry, Heinze and Gehlen,
2013). The skill of biological impact models, however,
needs to be tested and improved, and in many cases, we
may run up against the problem of predictability of complex
biological systems, especially as stakeholders ask more
focused questions related to individual species and specific
locations.
Finally, climate change and other human activities,
especially fishing and coastal habitat degradation, are neg-
atively impacting the marine resources and fisheries upon
which humans depend for food, personal security, and live-
lihoods (Allison et al., 2009; Cooley and Doney, 2009;
Halpern et al., 2012). More research is needed to evaluate
the efficacy of potential adaptation strategies and to better
understand the consequences for human, social, and eco-
nomic systems (Ruckelshaus et al., 2013).
ACKNOWLEDGMENTS
The author gratefully acknowledges support from the U.S. National
Science Foundation through the Palmer LTER project (http://pal.
lternet.edu/) (NSFOPP-0823101) and the Center forMicrobial Ocean-
ography Research and Education (C-MORE, http://cmore.soest.
hawaii.edu/) (NSF EF-0424599). The author thanks H. Ducklow,
K. Kroeker, C. Langdon, and D. Siegel for providing figures as well
as H. Ducklow, W. Gould, S. Sailley, O. Schofield, and J. Shepherd
for constructive comments.
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