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.

REFERENCES

Alheit, J., Bakun, A., 2010. Population synchronies within and between

ocean basins: apparent teleconnections and implications as to

physical–biological linkage mechanisms. J. Mar. Syst. 79, 267–285.

Allison, E.H., Perry, A.L., Badjeck, M.-C., Adger, W.N., Brown, K.,

Conway, D., Halls, A.S., Pilling, G.M., Reynolds, J.D.,

Andrew, N.L., Dulvy, N.K., 2009. Vulnerability of national economies

to the impacts of climate change on fisheries. Fish Fish. 10, 173–196.

Anderson, L.A., Sarmiento, J.L., 1994. Redfield ratios of remineralization

determined by nutrient data analysis. Global Biogeochem. Cycles 8,

65–80.

Anderson, D.M., Glibert, P.M., Burkholder, J.M., 2002. Harmful algal

blooms and eutrophication: nutrient sources, composition, and conse-

quences. Estuaries 25, 704–726.

Anderson, R.F.,Ali, S., Bradtmiller, L.I., Nielsen, S.H.H., Fleisher, M.Q.,

Anderson, B.E., Burckle, L.H., 2009. Wind-driven upwelling in the

Southern Ocean and the deglacial rise in atmospheric CO2. Science

323, 1443–1448.

Andersson, A.J., Mackenzie, F.T., Gattuso, J.-P., 2011. Effects of ocean

acidification on pelagic organism and ecosystems. In: Gattuso, J.P.,

Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 835

Hansson, L. (Eds.), Ocean Acidification. Oxford University Press,

Oxford, UK, pp. 122–153.

Anthony, K.R.N., Kline, D.I., Diaz-Pulido, G., Dove, S.,

Hoegh-Guldberg, O., 2008. Ocean acidification causes bleaching

and productivity loss in coral reef builders. Proc. Natl. Acad. Sci.

U.S.A. 105, 17442–17446.

Antoine, D., Morel, A., Gordon, H.R., Banzon, V.F., Evans, R.H., 2005.

Bridging ocean color observations of the 1980s and 2000s in search

of long-term trends. J. Geophys. Res. 110, C06009. http://dx.doi.

org/10.1029/2004JC002620.

Atkinson, A., Siegel, V., Pakhomov, E., Rothery, P., 2004. Long-term

decline in krill stock and increase in salps within the Southern Ocean.

Nature 432, 100–103.

Baker-Austin, C., Trinanes, J.A., Taylor, N.G.H., Hartnell, R.,

Siitonen, A., Martinez-Urtaza, J., 2013. Emerging Vibrio risk at

high latitudes in response to ocean warming. Nat. Clim. Chang. 3,

73–77.

Barton, A., Hales, B., Waldbusser, G.G., Langdon, C., Feely, R., 2012.

The Pacific oyster, Crassostrea gigas, shows negative correlation to

naturally elevated carbon dioxide levels: implications for near-term

ocean acidification effects. Limnol. Oceanogr. 57, 698–710.

Bates, N.R., Best, M.H.P., Neely, K., Garley, R., Dickson, A.G.,

Johnson, R.J., 2012. Detecting anthropogenic carbon dioxide uptake

and ocean acidification in the North Atlantic Ocean. Biogeosciences

9, 2509–2522.

Beaugrand, G., Reid, P.C., Ibanez, F., Lindley, J.A., Edwards, M., 2002.

Reorganization of North Atlantic marine copepod biodiversity and

climate. Science 296, 1692–1694.

Beaulieu, C., Henson, S.A., Sarmiento, J.L., Dunne, J.P., Doney, S.C.,

Rykaczewski, R.R., Bopp, L., 2013. Factors challenging our ability

to detect long-term trends in ocean chlorophyll. Biogeosciences 10,

2711–2724.

Behrenfeld, M.J., 2010. Abandoning Sverdrup’s critical depth hypothesis

on phytoplankton blooms. Ecology 91, 977–989.

Behrenfeld, M.J., Boss, E., Siegel, D.A., Shea, D.M., 2005. Carbon based

ocean productivity and phytoplankton physiology from space.

Global Biogeochem. Cycles 19, GB1006. http://dx.doi.org/10.1029/

2004GB002299.

Behrenfeld, M.J., O’Malley, R.T., Siegel, D.A., McClain, C.R.,

Sarmiento, J.L., Feldman, G.C., Milligan, A.J., Falkowski, P.G.,

Letelier, R.M., Boss, E.S., 2006. Climate-driven trends in contem-

porary ocean productivity. Nature 444, 752–755.

Beman, J.M., Chow, C.-E., King, A.L., Feng, Y., Fuhrman, J.A.,

Andersson, A., Bates, N.R., Popp, B.N., Hutchins, D.A., 2011. Global

declines in oceanic nitrification rates as consequence of ocean acidi-

fication. Proc. Natl. Acad. Sci. U.S.A. 108, 208–213.

Bender, M.A., Knutson, T.R., Tuleya, R.E., Sirutis, J.J., Vecchi, G.A.,

Garner, S.T., Held, I.M., 2010. Modeled impact of anthropogenic

warming on the frequency of intense Atlantic hurricanes. Science

327, 454–458.

Bernhardt, J.R., Leslie, H.M., 2013. Resilience to climate change in

coastal marine ecosystems. Ann. Rev. Mar. Sci. 5, 371–392.

Bestelmeyer,B.T.,Ellison,A.M.,Fraser,W.R.,Gorman,K.B.,Holbrook,S.J.,

Laney, C.M., Ohman, M.D., Peters, D.P.C., Pillsbury, F.C.,

Rassweiler, A., Schmitt, R.J., Sharma, S., 2011. Analysis of abrupt

transitions in ecological systems. Ecosphere 2 (12), 1–26, Article 129.

Bianchi, D., Dunne, J.P., Sarmiento, J.L., Galbraith, E.D., 2012. Data-

based estimates of suboxia, denitrification and N2O production

in the ocean and their sensitivities to dissolved oxygen. Global

Biogeochem. Cycles 26, GB2009. http://dx.doi.org/10.1029/

2011GB004209.

Bindoff, N.L., Willebrand, J., Artale, V., Cazenave, A., Gregory, J.,

Gulev, S., Hanawa, K., Le Quere, C., Levitus, S., Nojiri, Y.,

Shum, C.K., Talley, L.D., Unnikrishnan, A., 2007. Observations:

oceanic climate change and sea level. In: Solomon, S., Qin, D.,

Manning, M., Chen, Z., Marquis, M., Averyt, K.B., Tignor, M.,

Miller, H.L. (Eds.), Climate Change 2007: The Physical Science Basis,

Contribution of Working Group I to the Fourth Assessment Report of

the Intergovernmental Panel on Climate Change. Cambridge Uni-

versity Press, Cambridge, United Kingdom and New York, NY,

USA, pp. 385–432.

Bopp, L., Le Quere, C., Heimann, M., Manning, A.C., Monfray, P., 2002.

Climate induced oceanic oxygen fluxes: implications for the contem-

porary carbon budget. Global Biogeochem. Cycles 16, 1022. http://dx.

doi.org/10.1029/2001GB001445.

Bopp, L., Resplandy, L., Orr, J.C., Doney, S.C., Dunne, J.P., Gehlen, M.,

Halloran, P., Heinze, C., Ilyina, T., Seferian, R., Tjiputra, J.,

Vichi, M., 2013. Multiple stressors of ocean ecosystems in the 21st

century: projections with CMIP5 models. Biogeosci. Discuss. 10,

3627–3676.

Boyce, D.G., Lewis, M.R., Worm, B., 2010. Global phytoplankton decline

over the past century. Nature 466, 591–596.

Boyd, P.W., Doney, S.C., 2002. Modelling regional responses by marine

pelagic ecosystems to global climate change. Geophys. Res. Lett. 29

(16), 1806. http://dx.doi.org/10.1029/2001GL014130.

Boyd, P., Doney, S.C., 2003. The impact of climate change and feedback

process on the ocean carbon cycle. In: Fasham, M. (Ed.), Ocean Bio-

geochemistry. Springer, New York, NY, USA, pp. 157–193.

Boyd, P.W., Jickells, T., Law, C.S., Blain, S., Boyle, E.A.,

Buesseler, K.O., Coale, K.H., Cullen, J.J., de Baar, H.J.W.,

Follows, M., Harvey, M., Lancelot, C., Levasseur, M.,

Owens, N.P.J., Pollard, R., Rivkin, R.B., Sarmiento, J.,

Schoemann, V., Smetacek, V., Takeda, S., Tsuda, A., Turner, S.,

Watson, A.J., 2007. Mesoscale iron enrichment experiments 1993–

2005: synthesis and future directions. Science 315, 612–627.

Boyd, P.W., Doney, S.C., Strzepek, R., Dusenberry, J., Lindsay, K.,

Fung, I., 2008. Climate-mediated changes to mixed-layer properties

in the Southern Ocean: assessing the phytoplankton response. Bio-

geosciences 5, 847–864.

Brander, K.M., 2007. Global fish production and climate change. Proc.

Natl. Acad. Sci. U.S.A. 104, 19709–19714.

Brander, K., 2010. Impacts of climate change on fisheries. J. Mar. Syst.

79, 389–402.

Brewer, P.G., Peltzer, E.T., 2009. Limits to marine life. Science 324,

347–348.

Brinton, E., Townsend, A., 2003. Decadal variability in abundances of the

dominant euphausiid species in southern sectors of the California

Current. Deep Sea Res. Part II 50, 2469–2492.

Brodeur, R.D.,Ware, D.M., 1992. Interannual and interdecadal changes in

zooplankton biomass in the subarctic Pacific Ocean. Fish. Oceanogr.

1, 32–38.

Bryan, F.O., Danabasoglu, G., Nakashiki, N., Yoshida, Y., Kim, D.H.,

Tsutsui, J., Doney, S.C., 2006. Response of North Atlantic thermo-

haline circulation and ventilation to increasing carbon dioxide in

CCSM3. J. Clim. 19, 2382–2397.

Byrne, R.H., Mecking, S., Feely, R.A., Liu, X.W., 2010. Direct observa-

tions of basin-wide acidification of the North Pacific Ocean. Geophys.

Res. Lett. 37, L02601. http://dx.doi.org/10.1029/2009GL040999.

PART VI The Changing Ocean836

Cai, W.-J., Xinping Hu, X., Huang, W.-J., Murrell, M.C., Lehrter, J.C.,

Lohrenz, S.E., Chou, W.-C., Zhai, W., Hollibaugh, J.T., Wang, Y.,

Zhao, P., Guo, X., Gundersen, K., Dai, M., Gong, G.-C., 2011. Acid-

ification of subsurface coastal waters enhanced by eutrophication.

Nat. Geosci. 4, 766–770.

Cameron-Smith, P., Elliott, S., Maltrud, M., Erickson, D., Wingenter, O.,

2011. Changes in dimethyl sulfide oceanic distribution due to climate

change. Geophys. Res. Lett. 38, L07704. http://dx.doi.org/10.1029/

2011GL047069.

Chan, F., Barth, J.A., Lubchenco, J., Kirincich, A., Weeks, H.,

Peterson, W.T., Menge, B.A., 2008. Emergence of anoxia in the

California Current large marine ecosystem. Science 319, 920.

Charlson, R.J., Lovelock, J.E., Andreae, M.O., Warren, S.G., 1987.

Oceanic phytoplankton, atmospheric sulphur, cloud albedo and

climate. Nature 326, 655–661.

Chavez, F.P., Strutton, P.G., Friederich, C.E., Feely, R.A., Feldman, G.C.,

Foley, D.C., McPhaden, M.J., 1999. Biological and chemical response

of the equatorial Pacific Ocean to the 1997–98 El Nino. Science 286,

2126–2131.

Chavez, F.P., Ryan, J., Lluch-Cota, S.E., Niquen, M., 2003. From

anchovies to sardines and back: multidecadal change in the Pacific

Ocean. Science 299, 217–221.

Chavez, F.P., Messie, M., Pennington, J.T., 2011. Marine primary pro-

duction in relation to climate variability and change. Ann. Rev.

Mar. Sci. 3, 227–260.

Cheung, W.W.L., Lam, V.W.Y., Sarmiento, J.L., Kearney, K.,

Watson, R., Pauly, D., 2009. Projecting global marine biodiversity

impacts under climate change scenarios. Fish Fish. 10, 235–251.

Chisholm, S.W., 2000. Oceanography: stirring time in the Southern

Ocean. Nature 407, 685–687.

Codispoti, L.A., 2010. Interesting tines for marine N2O. Science 327,

1339–1340.

Cook, T., Folli, M., Klinck, J., Ford, S., Miller, J., 1998. The relationship

between increasing sea-surface temperature and the northward spread

of Perkinsus marinus (Dermo) disease epizootics in oysters. Estuar.

Coast. Shelf Sci. 46, 587–597.

Cooley, S.R., Doney, S.C., 2009. Anticipating ocean acidification’s eco-

nomic consequences for commercial fisheries. Environ. Res. Lett. 4,

024007. http://dx.doi.org/10.1088/1748-9326/4/2/024007.

Cooley, S., Kite-Powell, H.L., Doney, S.C., 2009. Ocean acidification’s

potential to alter global marine ecosystem services. Oceanography

22 (4), 172–180.

Cury, P., Bakun, A., Crawford, R.J.M., Jarre, A., Quinones, R.A.,

Shannon, L.J., Verheye, H.M., 2000. Small pelagics in upwelling

systems: patterns of interaction and structural changes in “wasp-waist”

ecosystems. ICES J. Mar. Sci. 57, 603–618.

Cushing, D.H., 1990. Plankton production and year-class strength in fish

populations: an update of the match/mismatch hypothesis. Adv. Mar.

Biol. 26, 249–293.

Dam, H.G., 2013. Evolutionary adaptation of marine zooplankton to

global change. Ann. Rev. Mar. Sci. 5, 349–370.

Davidson, E., 2009. The contribution of manure and fertilizer nitrogen to

atmospheric nitrous oxide since 1860. Nat. Geosci. 2, 659–662.

Denman, K.L., Brasseur, G., Chidthaisong, A., Ciais, P., Cox, P.M.,

Dickinson, R.E., Hauglustaine, D., Heinze, C., Holland, E., Jacob, D.,

Lohmann, U., Ramachandran, S., da Silva Dias, P.L., Wofsy, S.C.,

Zhang, X., 2007. Couplings between changes in the climate system

and biogeochemistry. In: Solomon, S., Qin, D., Manning, M.,

Chen, Z., Marquis, M., Averyt, K.B., Tignor, M., Miller, H.L. (Eds.),

Climate Change 2007: The Physical Science Basis, Contribution of

Working Group I to the Fourth Assessment Report of the Intergovern-

mental Panel on Climate Change. Cambridge University Press,

Cambridge, United Kingdom and New York, NY, USA, pp. 499–587.

Deutsch, C., Weber, T., 2012. Nutrient ratios as a tracer and driver of

ocean biogeochemistry. Ann. Rev. Mar. Sci. 4, 113–141.

Diaz, R.J., Rosenberg, R., 2008. Spreading dead zones and consequences

for marine ecosystems. Science 321, 926–929.

Doney, S.C., 1999. Major challenges confronting marine biogeochemical

modeling. Global Biogeochem. Cycles 13, 705–714.

Doney, S.C., 2006. Oceanography: plankton in a warmer world. Nature

444, 695–696.

Doney, S.C., 2010. The growing human footprint on coastal and open-

ocean biogeochemistry. Science 328, 1512–1516.

Doney, S.C., Sailley, S.F., 2013. When an ecological regime shift is really

just stochastic noise. Proc. Natl. Acad. Sci. U.S.A. 110 (7),

2438–2439.

Doney, S.C., Schimel, D.S., 2007. Carbon and climate system coupling on

timescales from the Precambrian to the Anthropocene. Ann. Rev.

Environ. Resour. 32, 31–66.

Doney, S.C., Fabry, V.J., Feely, R.A., Kleypas, J.A., 2009. Ocean acidifi-

cation: the other CO2 problem. Ann. Rev. Mar. Sci. 1, 169–192.

Doney, S.C., Ruckelshaus, M., Duffy, J.E., Barry, J.P., Chan, F.,

English, C.A., Galindo, H.M., Grebmeier, J.M., Hollowed, A.B.,

Knowlton, N., Polovina, J., Rabalais, N.N., Sydeman, W.J.,

Talley, L.D., 2012. Climate change impacts on marine ecosystems.

Ann. Rev. Mar. Sci. 4, 11–37.

Donner, S.D., 2009. Coping with commitment: projected thermal stress on

coral reefs under different future scenarios. PLoS One 4, e5712.

Dore, J.E., Popp, B.N., Karl, D.M., Sansone, F.J., 1998. A large source of

atmospheric nitrous oxide from subtropical North Pacific surface

waters. Nature 396, 63–66.

Dore, J.E., Lukas, R., Sadler, D.W., Church, M.J., Karl, D.M., 2009.

Physical and biogeochemical modulation of ocean acidification in

the central North Pacific. Proc. Natl. Acad. Sci. U.S.A. 106,

12235–12240.

Ducklow, H.W., Baker, K., Martinson, D.G., Quetin, L.B., Ross, R.M.,

Smith, R.C., Stammerjohn, S.E., Vernet, M., Fraser, W., 2007. Marine

pelagic ecosystems: the West Antarctic Peninsula. Philos. Trans. R.

Soc. B Biol. Sci. 362, 67–94.

Ducklow, H., Doney, S.C., Steinberg, D.K., 2009. Contributions of long-

term research and time-series observations to marine ecology and bio-

geochemistry. Ann. Rev. Mar. Sci. 1, 279–302.

Dunne, J.P., Sarmiento, J.L., Gnanadesikan, A., 2007. A synthesis of

global particle export from the surface ocean and cycling through

the ocean interior and on the seafloor. Global Biogeochem. Cycles

21, GB4006. http://dx.doi.org/10.1029/2006GB002907.

Eakin, C.M., Morgan, J.A., Heron, S.F., Smith, T.B., Liu, G.,

Alvarez-Filip, L., Baca, B., Bartels, E., Bastidas, C., Bouchon, C.,

2010. Caribbean corals in crisis: record thermal stress, bleaching,

and mortality in 2005. PLoS One 5, e13969.

Edwards,M., Richardson, A.J., 2004. Impact of climate change on marine

pelagic phenology and trophic mismatch. Nature 430, 881–884.

EPICA community members, 2004. Eight glacial cycles from an Ant-

arctic ice core. Nature 429, 623–628.

Eppley, R., 1972. Temperature and phytoplankton growth in the sea. Fish.

Bull. 70, 1063–1085.

Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 837

Evans, G.T., Parslow, J.S., 1985. A model of annual plankton cycles. Biol.

Oceanogr. 3, 327–347.

Fabricius, K.E., Langdon, C., Uthicke, S., Humphrey, C., Noonan, S.,

De’ath, G., Okazaki, R., Muehllehner, N., Glas, M.S., Lough, J.M.,

2011. Losers and winners in coral reefs acclimatized to elevated

carbon dioxide concentrations. Nat. Clim. Chang. 1, 165–169.

Fabry, V.J., Seibel, B.A., Feely, R.A., Orr, J.C., 2008. Impacts of ocean

acidification on marine fauna and ecosystem processes. ICES J.

Mar. Sci. 65, 414–432.

Feely, R.A., Sabine, C.L., Hernandez-Ayon, J.M., Ianson, D., Hales, B.,

2008. Evidence for upwelling of corrosive “acidified” water onto

the continental shelf. Science 320, 1490–1492.

Feely, R.A., Doney, S.C., Cooley, S.R., 2009. Ocean acidification: present

conditions and future changes in a high-CO2 world. Oceanography 22

(4), 36–47.

Feely, R.A., Alin, S.R., Newton, J., Sabine, C.L., Warner, M., Devol, A.,

Krembs, C., Maloy, C., 2010. The combined effects of ocean acidifi-

cation, mixing, and respiration on pH and carbonate saturation in an

urbanized estuary. Estuar. Coast. Shelf Sci. 88, 442–449.

Food and Agricultural Organization of the United Nations, 2007. The

world’s mangroves 1980–2005. FAO Forestry Paper 153. Food and

Agricultural Organization of the United Nations, Rome, Italy, 77 pp.

Food and Agricultural Organization of the United Nations, 2010.

Fishery and Aquaculture Statistics. FAO Yearbook 2008. Food and

Agricultural Organization of the United Nations, Rome, Italy, 72 pp.

Fox, D., 2012. Trouble bares its claws. Nature 492, 170–172.

Francis, T.B., Scheuerell, M.D., Brodeur, R.D., Levin, P.S., Ruzicka, J.J.,

Tolimieri, N., Peterson, W.T., 2012. Climate shifts the interaction web

of a marine plankton community. Glob. Chang. Biol 18, 2498–2508.

Freing, A., Wallace, D.W.R., Bange, H.W., 2012. Global oceanic pro-

duction of nitrous oxide. Philos. Trans. R. Soc. B 367, 1245–1255.

Friedlingstein, P., Cox, P., Betts, R., Bopp, L., von Bloh, W., Brovkin, V.,

Cadule, P., Doney, S., Eby, M., Fung, I., Bala, G., John, J., Jones, C.,

Joos, F., Kato, T., Kawamiya, M., Knorr, W., Lindsay, K.,

Matthews, H.D., Raddatz, T., Rayner, P., Reick, C., Roeckner, E.,

Schnitzler, K.-G., Schnur, R., Strassmann, K., Weaver, A.J.,

Yoshikawa, C., Zeng, N., 2006. Climate–carbon cycle feedback

analysis: results from the C4MIP model intercomparison. J. Clim.

19 (14), 3337–3353.

Frolicher, T.L., Joos, F., Plattner, G.-K., Steinacher, M., Doney, S.C.,

2009. Natural variability and anthropogenic trends in oceanic

oxygen in a coupled carbon cycle-climate model ensemble. Global

Biogeochem. Cycles 23, GB1003. http://dx.doi.org/10.1029/

2008GB003316.

Fung, I., Doney, S.C., Lindsay, K., John, J., 2005. Evolution of carbon

sinks in a changing climate. Proc. Natl. Acad. Sci. U.S.A. 102,

11201–11206.

Gardner, T.A., Cote, I.M., Gill, J.A., Grant, A., Watkinson, A.R., 2003.

Long-term region-wide declines in Caribbean corals. Science 301,

958–960.

Gattuso, J.P., Hansson, L. (Eds.), 2011. Ocean Acidification. Oxford

University Press, Oxford, UK, 326 pp.

Gattuso, J.P., Bijma, J., Gehlen, M., Riebesell, U., Turley, C., 2011.

Ocean acidification: knowns, unknowns, and perspectives. In:

Gattuso, J.P., Hansson, L. (Eds.), Ocean Acidification. Oxford

University Press, Oxford, UK, pp. 291–311.

Gehlen, M., Gruber, N., Gangst�, R., Bopp, L., Oschlies, A., 2011. Bio-

geochemical consequences of ocean acidification and feedbacks to

the earth system. In: Gattuso, J.P., Hansson, L. (Eds.), Ocean Acidifi-

cation. Oxford University Press, Oxford, UK, pp. 230–248.

Geider, R.J., LaRoche, J., 2002. Redfield revisited: variability of C:N:P

in marine microalgae and its biochemical basis. Eur. J. Phycol. 37,

1–17.

Geider, R.J.,MacIntyre, H.L., Kana, T.M., 1997. Dynamic model of phy-

toplankton growth and acclimation: responses of the balanced growth

rate and the chlorophyll a:carbon ratio to light, nutrient-limitation and

temperature. Mar. Ecol. Prog. Ser. 148, 187–200.

Glover, D.M., Jenkins, W.J., Doney, S.C., 2011. Modeling Methods for

Marine Science. Cambridge University Press, Cambridge, UK,

592 pp.

Gnanadesikan, A., Emanuel, K., Vecchi, G.A., Anderson, W.G.,

Hallberg, R., 2010. How ocean color can steer Pacific tropical

cyclones. Geophys. Res. Lett. 37, L18802. http://dx.doi.org/10.1029/

2010GL044514.

Grebmeier, J.M., 2012. Shifting patterns of life in the pacific Arctic and

sub-Arctic seas. Ann. Rev. Mar. Sci. 4, 63–78.

Grebmeier, J.M., Cooper, L.W., Feder, H.M., Sirenko, B.I., 2006. Eco-

system dynamics of the Pacific-influenced Northern Bering and

Chukchi Seas in the Amerasian Arctic. Prog. Oceanogr. 71, 331–361.

Grebmeier, J.M., Moore, S.E., Overland, J.E., Frey, K.E., Gradinger, R.,

2010. Biological response to recent pacific Arctic sea ice retreats. Eos

91, 161–163.

Griffis, R., Howard J. (Eds.), 2012. Oceans and Marine Resources in a

Changing Climate: Technical Input to the 2013 National

Climate Assessment. United States Global Change Research Project,

Washington, DC, 313 pp. Available from: http://downloads.usgcrp.

gov/NCA/technicalinputreports/Griffis_Howard_Ocean_Marine_Re-

sources.pdf

Gruber, N., 2011. Warming up, turning sour, losing breath: ocean biogeo-

chemistry under global change. Philos. Trans. R. Soc. A 369,

1980–1996.

Gruber, N., Kortzinger, A., Borges, A., Claustre, H., Doney, S.C.,

Feely, R.A., Hood, M., Ishii, M., Kozyr, A., Monteiro, P., Nojiri, Y.,

Sabine, C.L., Schuster, U., Wallace, D.W.R., Wanninkhof, R., 2010.

Toward an integrated observing system for ocean carbon and biogeo-

chemistry at a time of change. In: Hall, J., Harrison, D.E.,

Stammer, D. (Eds.), Proceedings of OceanObs’09: Sustained Ocean

Observations and Information for Society (Vol. 1). ESA Publication

WPP-306. http://dx.doi.org/10.5270/OceanObs09.pp.18.

Gruber, N., Hauri, C., Lachkar, Z., Loher, D., Frolicher, T.L.,

Plattner, G.-K., 2012. Rapid progression of ocean acidification in

the California current system. Science 337, 220–223.

Halloran, P.R., Bell, T.G., Totterdell, I.J., 2010. Can we trust empirical

marine DMS parameterisations within projections of future climate?

Biogeosciences 7, 1645–1656.

Hall-Spencer, J.M., Rodolfo-Metalpa, R., Martin, S., Ransome, E.,

Fine, M., Turner, S.M., Rowley, S.J., Tedesco, D., Buia, M.-C.,

2008. Volcanic carbon dioxide vents show ecosystem effects of ocean

acidification. Nature 454, 96–99.

Halpern, B.S., Longo, C., Hardy, D., McLeod, K.L., Samhouri, J.F.,

Katona, S.K., Kleisner, K., Lester, S.E., O’Leary, J., Ranelletti, M.,

Rosenberg, A.A., Scarborough, C., Selig, E.R., Best, B.D.,

Brumbaugh, D.R., Chapin, F.S., Crowder, L.B., Daly, K.L.,

Doney, S.C., Elfes, C., Fogarty, M.J., Gaines, S.D., Jacobsen, K.I.,

Karrer, L.B., Leslie, H.M., Neeley, E., Pauly, D., Polasky, S.,

Ris, B., St. Martin, K., Stone, G.S., Sumaila, U.R., Zeller, D., 2012.

PART VI The Changing Ocean838

An index to assess the health and benefits of the global ocean. Nature

488, 615–620.

Hare, S.R., Mantua, N.J., 2000. Empirical evidence for North Pacific

regime shifts in 1977 and 1989. Prog. Oceanogr. 47, 103–145.

Harvell, C.D., Mitchell, C.E., Ward, J.R., Altizer, S., Dobson, A.P.,

Ostfeld, R.S., Samuel, M.D., 2002. Climate warming and disease risks

for terrestrial and marine biota. Science 296, 2158–2162.

Heinze, C., Gehlen, M., 2013. Modeling ocean biogeochemical processes

and resulting tracer distributions. In: Siedler, G., Church, J., Gould, J.,

Griffies, S. (Eds.), Ocean Circulation and Climate. second ed.

Academic Press (Elsevier), Waltham, MA, USA.

Helm, K.P., Bindoff, N.L., Church, J.A., 2011. Observed decreases in

oxygen content of the global ocean. Geophys. Res. Lett. 38,

L23602. http://dx.doi.org/10.1029/2011GL049513.

Henson, S.A., Dunne, J.P., Sarmiento, J.L., 2009. Decadal variability in

North Atlantic phytoplankton blooms. J. Geophys. Res. 114,

C04013. http://dx.doi.org/10.1029/2008JC005139.

Henson, S.A., Sarmiento, J.L., Dunne, J.P., Bopp, L., Lima, I.,

Doney, S.C., John, J., Beaulieu, C., 2010. Detection of anthropogenic

climate change in satellite records of ocean chlorophyll and produc-

tivity. Biogeosciences 7, 621–640.

Henson, S.A., Sanders, R., Madsen, E., Morris, P.J., Le Moigne, F.,

Quartly, G.D., 2011. A reduced estimate of the strength of the ocean’s

biological carbon pump. Geophys. Res. Lett. 38, L04606. http://dx.doi.

org/10.1029/2011GL046735.

Hoegh-Guldberg, O., Mumby, P.J., Hooten, A.J., Steneck, R.S.,

Greenfield, P., Gomez, E., Harvell, C.D., Sale, P.F., Edwards, A.J.,

Caldeira, K., Knowlton, N., Eakin, C.M., Iglesias-Prieto, R.,

Muthiga, N., Bradbury, R.H., Dubi, A., Hatziolos, M.E., 2007. Coral

reefs under rapid climate change and ocean acidification. Science 318,

1737–1742.

Hofmann, G.E., Smith, J.E., Johnson, K.S., Send, U., Levin, L.A.,

Micheli, F., Paytan, A., Price, N.M., Peterson, B., Takeshita, Y.,

Matson, P.G., Crook, E.D., Kroeker, K.J., Gambi, M.C.,

Rivest, E.B., Frieder, C.A., Yu, P.C., Martz, T.R., 2011. High-

frequency dynamics of ocean pH: a multi-ecosystem comparison.

PLoS One 6 (12), e28983.

Hollowed, A.B., Hare, S.R., Wooster, W.S., 2001. Pacific basin climate

variability and patterns of Northeast Pacific marine fish production.

Prog. Oceanogr. 49, 257–282.

Honisch, B., Ridgwell, A., Schmidt, D.N., Thomas, E., Gibbs, S.J.,

Sluijs, A., Zeebe, R., Kump, L., Martindale, R.C., Greene, S.E.,

Kiessling, W., Ries, J., Zachos, J.C., Royer, D.L., Barker, S.,

Marchitto Jr., T.M., Moyer, R., Pelejero, C., Ziveri, P., Foster, G.L.,

Williams, B., 2012. The geological record of ocean acidification.

Science 335, 1058–1063.

Hsieh, C.H., Glaser, S.M., Lucas, A.J., Sugihara, G., 2005. Distinguishing

random environmental fluctuations from ecological catastrophes for

the North Pacific Ocean. Nature 435, 336–340.

Hunt Jr., G.L., Coyle, K.O., Eisner, L.B., Farley, E.V., Heintz, R.A.,

Mueter, F., Napp, J.M., Overland, J.E., Ressler, P.H., Salo, S.,

Stabeno, P.J., 2011. Climate impacts on eastern Bering Sea foodwebs:

a synthesis of new data and an assessment of the Oscillating Control

Hypothesis. ICES J. Mar. Sci. 68, 1230.

Ianelli, J.N., Hollowed, A.B., Haynie, A.C., Mueter, F.J., Bond, N.A.,

2011. Evaluating management strategies for eastern Bering Sea

walleye pollock (Theragra chalcogramma) in a changing envi-

ronment. ICES J. Mar. Sci. 68, 1297–1304.

Iglesias-Rodriguez, M.D., Fabry, V.J., Dickson, A., Gattuso, J.-P.,

Bijma, J., Riebesell, U., Doney, S., Turley, C., Saino, T., Lee, K.,

Anthony, K., Kleypas, J., 2010. Developing a global ocean acidifi-

cation observation network. In: Hall, J., Harrison, D.E.,

Stammer, D. (Eds.), Proceedings of OceanObs’09: Sustained Ocean

Observations and Information for Society. vol. 1. ESA Publication

WPP-306. http://dx.doi.org/10.5270/OceanObs09.pp.24.

Ito, T., Follows, M.J., 2005. Preformed phosphate, soft tissue pump and

atmospheric CO2. J. Mar. Res. 63 (4), 813–839.

Jickells, T.D., An, Z.S., Andersen, K.K., Baker, A.R., Bergametti, G.,

Brooks, N., Cao, J.J., Boyd, P.W., Duce, R.A., Hunter, K.A.,

Kawahata, H., Kubilay, N., LaRoche, J., Liss, P.S., Mahowald, N.,

Prospero, J.M., Ridgwell, A.J., Tegen, I., Torres, R., 2005. Global iron

connections between desert dust, ocean biogeochemistry, and climate.

Science 308, 67–71.

Jochum, M., Yeager, S., Lindsay, K., Moore, K., Murtugudde, R., 2010.

Quantification of the feedback between phytoplankton and ENSO in

the Community Climate System Model. J. Clim. 11, 2916–2925.

Johnson, K.S., Berelson, W.M., Boss, E.S., Chase, Z., Claustre, H.,

Emerson, S.R., Gruber, N., Kortzinger, A., Perry, M.J., Riser, S.C.,

2009. Observing biogeochemical cycles at global scales with profiling

floats and gliders: prospects for a global array. Oceanography 22 (3),

216–225.

Jones, A., Berkelmans, R., 2010. Potential costs of acclimatization to a

warmer climate: growth of a reef coral with heat tolerant versus sen-

sitive symbiont types. PLoS One 5, e10437.

Karl, D., Letelier, R., Tupas, L., Dore, J., Christian, J., Hebel, D., 1997.

The role of nitrogen fixation in biogeochemical cycling in the sub-

tropical North Pacific Ocean. Nature 388, 533–538.

Kawasaki, T., 1992. Mechanisms governing fluctuations in pelagic fish

populations. S. Afr. J. Mar. Sci. 12, 873–879.

Keeling, R.F.,Kortzinger, A., Gruber, N., 2010. Ocean deoxygenation in a

warming world. Ann. Rev. Mar. Sci. 2, 199–229.

Khatiwala, S., Primeau, F., Hall, T., 2009. Reconstruction of the history of

anthropogenic CO2 concentrations in the ocean. Nature 462, 346–349.

Kiehl, J.T., Trenberth, K.E., 1997. Earth’s annual global mean energy

budget. Bull. Am. Meteorol. Soc. 78 (2), 197–208.

Kline, D.I., Teneva, L., Schneider, K., Miard, T., Chai, A., Marker, M.,

Headley, K., Opdyke, B., Nash, M., Valetich, M., Caves, J.K.,

Russell, B.D., Connell, S.D., Kirkwood, B.J., Brewer, P.,

Peltzer, E., Silverman, J., Caldeira, K., Dunbar, R.B., Koseff, J.R.,

Monismith, S.G., Mitchell, B.G., Dove, S., Hoegh-Guldberg, O.,

2012. A short-term in situ CO2 enrichment experiment onHeron Island

(GBR). Sci. Rep. 2, 413. http://dx.doi.org/10.1038/srep00413.

Kroeker, K.J., Kordas, R.L., Crim, R.N., Singh, G.G., 2010. Meta-

analysis reveals negative yet variable effects of ocean acidification

on marine organisms. Ecol. Lett. 13, 1419–1434.

Kroeker, K.J., Kordas, R.L., Crim, R.N., Hendriks, I.E., Ramajo, L.,

Singh, G.G., Duarte, C., Gattuso, J.P., 2013. Impacts of ocean acidifi-

cation on marine organisms: quantifying sensitivities and interaction

with warming. Glob. Chang. Biol 19, 1884–1896. http://dx.doi.org/

10.1111/gcb.12179.

Kwon, E.Y., Primeau, F., Sarmiento, J.L., 2009. The impact of reminera-

lization depth on the air-sea carbon balance. Nat. Geosci. 2, 630–635.

Lavaniegos, B.,Ohman, M., 2003. Long-term changes in pelagic tunicates

of the California Current. Deep Sea Res. Part II 50, 2473–2498.

Le Quere, C., Rodenbeck, C., Buitenhuis, E.T., Conway, T.J.,

Langenfelds, R., Gomez, A., Labuschagne, C., Ramonet, M.,

Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 839

Nakazawa, T., Metzl, N., Gillett, N., Heimann, M., 2007. Saturation of

the southern ocean CO2 sink due to recent climate change. Science

316, 1735–1738.

Le Quere, C., Raupach, M.R., Canadell, J.G., Marland, G., Bopp, L.,

Ciais, P., Conway, T.J., Doney, S.C., Feely, R.A., Foster, P.,

Friedlingstein, P., Gurney, K., Houghton, R.A., House, J.I.,

Huntingford, C., Levy, P.E., Lomas, M.R., Majkut, J., Metzl, N.,

Ometto, J.P., Peters, G.P., Prentice, I.C., Randerson, J.T.,

Running, S.W., Sarmiento, J.L., Schuster, U., Sitch, S.,

Takahashi, T., Viovy, N., van der Werf, G.R., Woodward, F.I.,

2009. Trends in the sources and sinks carbon dioxide. Nat. Geosci.

2, 831–836.

Le Quere, C., Takahashi, T., Buitenhuis, E.T., Rodenbeck, C.,

Sutherland, S.C., 2010. Impact of climate change and variability on

the global oceanic sink of CO2. Global Biogeochem. Cycles 24,

GB4007. http://dx.doi.org/10.1029/2009GB003599.

Levin, L.A., Ekau, W., Gooday, A.J., Jorissen, F., Middelburg, J.J.,

Naqvi, S.W.A., Neira, C., Rabalais, N.N., Zhang, J., 2009. Effects

of natural and human-induced hypoxia on coastal benthos. Biogeos-

ciences 6, 2063–2098.

Loeng, H., Brander, K., Carmack, E., Denisenko, S., Drinkwater, K.,

Hansen, B., Kovacs, K., Livingston, P., McLaughlin, F.,

Saksburg, E., 2005. Marine systems. In: Symon, C., Arris, L.,

Heal, B. (Eds.), Arctic Climate Impact Assessment. Cambridge

University Press, Cambridge, UK, pp. 453–538.

Lovenduski, N., Gruber, N., Doney, S.C., 2008. Towards a mechanistic

understanding of the decadal trends in the Southern Ocean carbon sink.

Global Biogeochem. Cycles 22, GB3016. http://dx.doi.org/10.1029/

2007GB003139.

MacFarling Meure, C., Etheridge, D., Trudinger, C., Steele, P.,

Langenfelds, R., van Ommen, T., Smith, A., Elkins, J., 2006. The

LawDome CO2, CH4 and N2O ice core records extended to 2000 years

BP. Geophys. Res. Lett. 33, L14810. http://dx.doi.org/10.1029/

2006GL026152.

Mackas, D.L., Beaugrand, G., 2010. Comparisons of zooplankton time

series. J. Mar. Syst. 79, 286–304.

Mann, K.H., Lazier, J.R., 2005. Dynamics of Marine Ecosystems: Bio-

logical–Physical Interactions in the Oceans. Wiley & Blackwell,

Malden, Massachusetts, 512 pp.

Mantua, N., Hare, S., Zhang, Y., Wallace, J., Francis, R., 1997. A Pacific

interdecadal climate oscillation with impacts on salmon production.

Bull. Am. Meteorol. Soc. 78, 1069–1079.

Marinov, I., Gnanadesikan, A., Toggweiler, J.R., Sarmiento, J.L., 2006.

The southern ocean biogeochemical divide. Nature 441, 964–967.

Marinov, I., Doney, S.C., Lima, I.D., 2010. Response of ocean phyto-

plankton community structure to climate change over the 21st century:

partitioning the effects of nutrients, temperature and light. Biogeos-

ciences 7, 3941–3959.

Martin, J.H., 1990. Glacial-interglacial CO2 change: the iron hypothesis.

Paleoceanography 5, 1–13.

Martin, J.H., Fitzwater, S.E., 1988. Iron deficiency limits phytoplankton

growth in the north-east Pacific subarctic. Nature 331, 341–343.

Martiny, A.C., Pham, C.T.A., Primeau, F.W., Vrugt, J.A., Moore, J.K.,

Levin, S.A., Lomas, M.W., 2013. Strong latitudinal patterns in the ele-

mental ratios of marine plankton and organic matter. Nat. Geosci. 6,

279–283.

McClain, C.R., 2009. A decade of satellite ocean color observations.

Ann. Rev. Mar. Sci. 1, 19–42.

McGillicuddy Jr., D.J.,Anderson, L.A., Doney, S.C., Maltrud, M.E., 2003.

Eddy-driven sources and sinks of nutrients in the upper ocean: results

from a 0.1 resolution model of the North Atlantic. Global Biogeochem.

Cycles 17 (2), 1035. http://dx.doi.org/10.1029/2002GB001987.

Menviel, L., Joos, F., 2012. Toward explaining the Holocene carbon

dioxide and carbon isotope records: results from transient ocean

carbon cycle-climate simulations. Paleoceanography 27, PA1207.

http://dx.doi.org/10.1029/2011PA002224.

Miller, C.B., Wheeler, P.A., 2012. Biological Oceanography, second ed.

John Wiley & Sons, Oxford, UK, 480 pp.

Miller, J., Muller, E., Rogers, C.S., Waara, R., Atkinson, A.,

Whelan, K.R.T., Patterson, M., Witcher, B., 2009. Coral disease fol-

lowing massive bleaching in 2005 causes 60% decline in coral cover

on reefs in the US Virgin Islands. Coral Reefs 28, 925–937.

Montes-Hugo, M., Doney, S.C.,Ducklow,H.W., Fraser,W.,Martinson,D.,

Stammerjohn, S.E., Schofield, O., 2009. Recent changes in phyto-

plankton communities associated with rapid regional climate change

along the western Antarctic Peninsula. Science 323, 1470–1473.

Moore, J.K., Doney, S.C., Lindsay, K., Mahowald, N., Michaels, A.F.,

2006. Nitrogen fixation amplifies the ocean biogeochemical response

to decadal timescale variations in mineral dust deposition. Tellus 58B,

560–572.

Moran, X.A.G., Lopez-Urrutia, A., Calvo-Diaz, A., Li, W.K.W., 2010.

Increasing importance of small phytoplankton in a warmer ocean.

Glob. Chang. Biol 16 (3), 1137–1144.

Morel, A., Antoine, D., 1994. Heating rate within the upper ocean in

relation to its biooptical state. J. Phys. Oceanogr. 24, 1652–1665.

Mueter, F.J., Bond, N.A., Ianelli, J.N., Hollowed, A.B., 2011. Expected

declines in recruitment of walleye pollock (Theragra chalcogramma)

in the eastern Bering Sea under future climate change. ICES J. Mar.

Sci. 68, 1284–1296.

Murtugudde, R., Beauchamp, J., McClain, C.R., Lewis, M.,

Busalacchi, A.J., 2002. Effects of penetrative radiation on the upper

tropical ocean circulation. J. Clim. 15, 470–486.

Nevison, C., Butler, J.H., Elkins, J.W., 2003. Global distribution of N2O

and DN2O–AOU yield in the subsurface ocean. Global Biogeochem.

Cycles 17, 1119, http://dx.doi.org/10.029/2003GB002068.

Nye, J.A., Link, J.S., Hare, J.A., Overholtz, W.J., 2009. Changing spatial

distribution of fish stocks in relation to climate and population size on

the Northeast United States continental shelf. Mar. Ecol. Prog. Ser.

393, 111–129.

Ohlmann, J.C., Siegel, D.A., Gautier, C., 1996. Ocean mixed layer radiant

heating and solar penetration: a global analysis. J. Clim. 9, 2265–2280.

Orr, J.C., Fabry, V.J., Aumont, O., Bopp, L., Doney, S.C., Feely, R.A.,

Gnanadesikan, A., Gruber, N., Ishida, A., Joos, F., Key, R.M.,

Lindsay, K., Maier-Reimer, E., Matear, R., Monfray, P.,

Mouchet, A., Najjar, R.G., Plattner, G.-K., Rodgers, K.B.,

Sabine, C.L., Sarmiento, J.L., Schlitzer, R., Slater, R.D.,

Totterdell, I.J., Weirig, M.-F., Yamanaka, Y., Yool, A., 2005. Anthro-

pogenic ocean acidification over the twenty-first century and its

impact on marine calcifying organisms. Nature 437, 681–686.

Oschlies, A., Shulz, K.G., Riebesell, U., Schmittner, A., 2008. Simulated

21st century’s increase in oceanic suboxia by CO2-enhanced biotic

carbon export. Global Biogeochem. Cycles 22, GB4008. http://dx.

doi.org/10.1029/2007GB003147.

Overland, J.E., Alheit, J., Bakun, A., Hurrell, J.W., Mackas, D.L.,

Miller, A.J., 2010. Climate controls on marine ecosystems and fish

populations. J. Mar. Syst. 79, 305–315.

PART VI The Changing Ocean840

Park, G.-H., Wanninkhof, R., Doney, S.C., Takahashi, T., Lee, K.,

Feely, R.A., Sabine, C.L., Trinanes, J., Lima, I.D., 2010. Variability

of global net air-sea CO2 fluxes over the last three decades using

empirical relationships. Tellus 62B, 352–368.

Parmesan, C., 2006. Ecological and evolutionary responses to recent

climate change. Ann. Rev. Ecol. Evol. Syst. 37, 637–669.

Paulmier, A.,Ruiz-Pino, D., 2009. Oxygenminimum zones (OMZs) in the

modern ocean. Prog. Oceanogr. 80, 113–128.

Perry, A.L., Low, P.J., Ellis, J.R., Reynolds, J.D., 2005. Climate change

and distribution shifts in marine fishes. Science 308, 1912–1915.

Polovina, J.J., Howell, E.A., Abecassis, M., 2008. Ocean’s least pro-

ductive waters are expanding. Geophys. Res. Lett. 35, L03618.

http://dx.doi.org/10.1029/2007GL031745.

Polovina, J.J., Dunne, J.P., Woodworth, P.A., Howell, E.A., 2011. Pro-

jected expansion of the subtropical biome and contraction of the tem-

perate and equatorial upwelling biomes in the North Pacific under

global warming. ICES J. Mar. Sci. 68, 986–995.

Portner, H.O., Farrell, A.P., 2008. Ecology: physiology and climate

change. Science 322, 690–692.

Portner, A.P., Gutowska, M., Ishimatsu, A., Lucassen, M., Melzner, F.,

Seibel, B., 2011. Effects of ocean acidification on nektonic organisms.

In: Gattuso, J.P., Hansson, L. (Eds.), Ocean Acidification. Oxford

University Press, Oxford, UK, pp. 154–175.

Quinn, P.K., Bates, T.S., 2011. The case against climate regulation via

oceanic phytoplankton sulphur emissions. Nature 480, 51–56.

Rabalais, N.N., Diaz, R.J., Levin, L.A., Turner, R.E., Gilbert, D.,

Zhang, J., 2010. Dynamics and distribution of natural and human-

caused hypoxia. Biogeosciences 7, 585–619.

Redfield, A.C., 1958. The biological control of chemical factors in the

environment. Am. Sci. 46, 205–221.

Redfield, A.C., Ketchum, B.H., Richards, F.A., 1963. The influence of

organisms on the composition of seawater. In: Hill, M.N. (Ed.), The

Sea, 2. Interscience, New York, USA, pp. 26–77.

Riebesell, U., Tortell, P.D., 2011. Effects of ocean acidification on pelagic

organism and ecosystems. In: Gattuso, J.P., Hansson, L. (Eds.), Ocean

Acidification. Oxford University Press, Oxford, UK, pp. 99–121.

Roemmich, D., McGowan, J., 1995. Climatic warming and the decline of

zooplankton in the California Current. Science 267, 1324–1326.

Royer, D.L., Pagani, M., Beerling, D.J., 2012. Geobiological constraints

on Earth system sensitivity to CO2 during the Cretaceous and

Cenozoic. Geobiology 10, 298–310.

Ruckelshaus, M., Doney, S.C., Galindo, H.M., Barry, J.P., Chan, F.,

Duffy, J.E., English, C.A., Gaines, S.D., Grebmeier, J.M.,

Hollowed, A.B., Knowlton, N., Polovina, J., Rabalais, N.N.,

Sydeman,W.J., Talley, L.D., 2013. Securing ocean benefits for society

in the face of climate change. Mar. Policy, 40, 154–159.

Rykaczewski, R.R., Checkley Jr., D.M., 2008. Influence of ocean winds

on the pelagic ecosystem in upwelling regions. Proc. Natl. Acad.

Sci. U.S.A. 105, 1965–1970.

Rykaczewski, R.R., Dunne, J.P., 2011. A measured look at ocean chloro-

phyll trends. Nature 472, E5–E6. http://dx.doi.org/10.1038/

nature09952.

Saba, V.S., Friedrichs, M.A.M., Carr, M.-E., Antoine, D.,

Armstrong, R.A., Asanuma, I., Aumont, O., Bates, N.R.,

Behrenfeld, M.J., Bennington, V., Bopp, L., Bruggeman, J.,

Buitenhuis, E.T., Church, M.J., Ciotti, A.M., Doney, S.C.,

Dowell, M., Dunne, J., Dutkiewicz, S., Gregg, W., Hoepffner, N.,

Hyde, K.J.W., Ishizaka, J., Kameda, T., Karl, D.M., Lima, I.,

Lomas, M.W., Marra, J., McKinley, G.A., Melin, F., Moore, J.K.,

Morel, A., O’Reilly, J., Salihoglu, B., Scardi, M., Smyth, T.J.,

Tang, S., Tjiputra, J., Uitz, J., Vichi, M., Waters, K.,

Westberry, T.K., Yool, A., 2010. Challenges of modeling depth-

integrated marine primary productivity over multiple decades: a case

study at BATS and HOT. Global Biogeochem. Cycles 24, GB3020.

http://dx.doi.org/10.1029/2009GB003655.

Sabine, C.L., Toste Tanhua, T., 2010. Estimation of anthropogenic CO2

inventories in the ocean. Ann. Rev. Mar. Sci. 2, 175–198.

Sabine, C.L., Feely, R.A., Gruber, N., Key, R.M., Lee, K.,

Gruber, Nicolas, Bullister, J.L., Wanninkhof, R., Wong, C.S.,

Wallace, D.W.R., Tilbrook, B., Millero, F.J., Peng, T.-H.,

Kozyr, A., Ono, T., Rios, A.F., 2004. The oceanic sink for anthropo-

genic CO2. Science 305, 367–371.

Sarmiento, J.L., Gruber, N., 2002. Sinks for anthropogenic carbon. Phys.

Today 55, 30–36.

Sarmiento, J.L., Gruber, N., 2006. Ocean Biogeochemical Dynamics.

Princeton University Press, Princeton, NJ, 503 pp.

Sarmiento, J.L., Orr, J.C., Siegenthaler, U., 1992. A perturbation simu-

lation of CO2 uptake in an ocean general circulation model. J.

Geophys. Res. 97, 3621–3645.

Sarmiento, J.L., Le Quere, C., 1996. Oceanic carbon dioxide uptake in a

model of century-scale global warming. Science 274, 1346–1350.

Sarmiento, J.L., Gruber, N., Brzezinski, M.A., Dunne, J.P., 2004a. High-

latitude controls of thermocline nutrients and low latitude biological

productivity. Nature 427, 56–60.

Sarmiento, J., Slater, R., Barber, R., Bopp, L., Doney, S.C., Hirst, A.C.,

Kleypas, J., Matear, R., Mikolajewicz, U., Monfray, P., Soldatov, V.,

Spall, S., Slater, R., Stouffer, R., 2004b. Response of ocean ecosystems

to climate warming. Global Biogeochem. Cycles 18, GB3003. http://

dx.doi.org/10.1029/2003GB002134.

Schofield, O., Ducklow, H.W., Martinson, D.G., Meredith, M.P.,

Moline, M.A., Fraser, W.R., 2010. How do polar marine ecosystems

respond to rapid climate change? Science 328, 1520–1523.

Seitzinger, S.P.,Mayorga,E.,Bouwman,A.F.,Kroeze,C.,Beusen,A.H.W.,

Billen, G., Van Drecht, G., Dumont, E., Fekete, B.M., Garnier, J.,

Harrison, J.A., 2010. Global river nutrient export: a scenario analysis

of past and future trends. Global Biogeochem. Cycles 24, GB0A08.

http://dx.doi.org/10.1029/2009GB003587.

Shakun, J.D., Clark, P.U., He, F., Marcott, S.A., Mix, A.C., Liu, Z.,

Otto-Bliesner, B., Schmittner, A., Bard, E., 2012a. Global warming

preceded by increasing carbon dioxide concentrations during the last

deglaciation. Nature 484, 49–54.

Shakun, J.D., Clark, P.U., He, F., Marcott, S.A., Mix, A.C., Liu, Z.,

Otto-Bliesner, B., Schmittner, A., Bard, E., 2012b. Global warming

preceded by increasing carbon dioxide concentrations during the last

deglaciation. Nature 484, 49–54.

Siegel, D.A., Doney, S.C., Yoder, J.A., 2002. The North Atlantic spring

phytoplankton bloom and Sverdrup’s critical depth hypothesis.

Science 296, 730–733.

Siegel, D.A., Behrenfeld, M.J., Maritorena, S., McClain, C.R., Antoine, D.,

Bailey, S.W., Bontempi, P.S., Boss, E.S., Dierssen, H.M., Doney, S.C.,

Eplee Jr., R.E., Evans, R.H., Feldman, G.C., Fields, E., Franz, B.A.,

Kuring, N.A., Mengelt, C., Nelson, N.B., Patt, F.S., Robinson, W.D.,

Sarmiento, J.L., Swan, C.S., Werdell, P.J., Westberry, T.K.,

Wilding, J.G., Yoder, J.A., 2013. Regional to global assessments of phy-

toplanktondynamics from the SeaWiFSmission.RemoteSens. Environ.

135, 77–91.

Chapter 31 Marine Ecosystems, Biogeochemistry, and Climate 841

Siegenthaler, U., Stocker, T.F., Monnin, E., Luthi, D., Schwander, J.,

Stauffer, B., Raynaud, D., Barnola, J.-M., Fischer, H.,

Masson-Delmotte, V., Jouzel, J., 2005. Stable carbon cycle-climate

relationship during the late Pleistocene. Science 310, 1313–1317.

Sigler, M.F., Renner, M., Danielson, S.L., Eisner, L.B., Lauth, R.R.,

Kuletz, K.J., Logerwell, E.A., Hunt Jr., G.L., 2011. Fluxes, fins, and

feathers: relationships among the Bering, Chukchi, and Beaufort Seas

in a time of climate change. Oceanography 24 (3), 250–265.

Sigman, D.M., Boyle, E.A., 2000. Glacial/interglacial variations in atmo-

spheric carbon dioxide. Nature 407, 859–869.

Smith, T.M., Reynolds, R.W., Peterson, T.C., Lawrimore, J., 2008.

Improvements to NOAA’s historical merged land-ocean surface tem-

perature analysis (1880–2006). J. Clim. 21, 2283–2296.

Smith, C.R., Grange, L.J., Honig, D.L., Naudts, L., Huber, B., Guidi, L.,

Domack, E., 2012. A large population of king crabs in the Palmer Deep

on the west Antarctic Peninsula shelf and potential invasive impacts.

Proc. R. Soc. B 279, 1017–1026.

Solomon, S., Qin, D., Manning, M., Alley, R.B., Berntsen, T.,

Bindoff, N.L., Chen, Z., Chidthaisong, A., Gregory, J.M.,

Hegerl, G.C., Heimann, M., Hewitson, B., Hoskins, B.J., Joos, F.,

Jouzel, J., Kattsov, V., Lohmann, U., Matsuno, T., Molina, M.,

Nicholls, N., Overpeck, J., Raga, G., Ramaswamy, V., Ren, J.,

Rusticucci, M., Somerville, R., Stocker, T.F., Whetton, P.,

Wood, R.A., Wratt, D., 2007. Technical summary. In: Solomon, S.,

Qin, D., Manning, M., Chen, Z., Marquis, M., Averyt, K.B.,

Tignor, M., Miller, H.L. (Eds.), Climate Change 2007: The Physical

Science Basis, Contribution of Working Group I to the Fourth

Assessment Report of the Intergovernmental Panel on Climate

Change. Cambridge University Press, Cambridge, United Kingdom

and New York, NY, USA, pp. 19–91.

Somero, G.N., 2012. The physiology of global change: linking patterns to

mechanisms. Ann. Rev. Mar. Sci. 4, 39–61.

Steinacher, M., Joos, F., Frolicher, T.L., Plattner, G.-K., Doney, S.C.,

2009. Imminent ocean acidification in the Arctic projected with the

NCAR global coupled carbon cycle-climate model. Biogeosciences

6, 515–533.

Steinacher, M., Joos, F., Frolicher, T.L., Bopp, L., Cadule, P., Cocco, V.,

Doney, S.C., Gehlen, M., Lindsay, K., Moore, J.K., Schneider, B.,

Segschneider, J., 2010. Projected 21st century decrease in marine pro-

ductivity: a multi-model analysis. Biogeosciences 7, 979–1005.

Steinberg, D.K., Lomas, M.W., Cope, J.S., 2012. Long-term increase in

mesozooplankton biomass in the Sargasso Sea: linkage to climate

and implications for food web dynamics and biogeochemical cycling.

Global Biogeochem. Cycles 26, GB1004. http://dx.doi.org/10.1029/

2010GB004026.

Stenseth, N.C., Mysterud, A., 2002. Climate, changing phenology, and

other life history traits: nonlinearity and match–mismatch to the envi-

ronment. Proc. Natl. Acad. Sci. U.S.A. 99 (21), 13379–13381.

Stenseth, N.C., Ottersen, G., Hurrell, J.W., Mysterud, A., Lima, M.,

Chan, K.-S., Yoccoz, N.G., Adlandsvik, B., 2003. Studying climate

effects on ecology through the use of climate indices: the North

Atlantic Oscillation, El Nino Southern Oscillation and beyond. Proc.

R. Soc. B Biol. Sci. 270 (1529), 2087–2096.

Stock, C.A., Alexander, M.A., Bond, N.A., Brander, K.M.,

Cheung, W.W.L., Curchitser, E.N., Delworth, T.L., Dunne, J.P.,

Griffies, S.M., Haltuch, M.A., Hare, J.A., Hollowed, A.B.,

Lehodey, P., Levin, S.A., Link, J.S., Rose, K.A., Rykaczewski, R.R.,

Sarmiento, J.L., Stouffer, R.J., Schwing, F.B., Vecchi, G.A.,

Werner, F.E., 2011. On the use of IPCC-class models to assess the

impact of climate on living marine resources. Prog. Oceanogr. 88, 1–27.

Stramma, L., Johnson, G.C., Sprintall, J., Mohrholz, V., 2008. Expanding

oxygen-minimum zones in the tropical oceans. Science 320, 655–658.

Stramma, L., Prince, E.D., Schmidtko, S., Jiangang Luo, J.,

Hoolihan, J.P., Visbeck, M., Wallace, D.W.R., Brandt, P.,

Kortzinger, A., 2012. Expansion of oxygen minimum zones may

reduce available habitat for tropical pelagic fishes. Nat. Clim. Chang.

2, 33–37.

Stroeve, J.C., Kattsov, V., Barrett, A., Serreze, M., Pavlova, T.,

Holland, M., Meier, W.N., 2012. Trends in Arctic sea ice extent from

CMIP5, CMIP3 and observations. Geophys. Res. Lett. 39, L16502.

http://dx.doi.org/10.1029/2012GL052676.

Sumaila, U.R., Cheung, W.W.L., Lam, V.W.Y., Pauly, D., Herrick, S.,

2011. Climate change impacts on the biophysics and economics of

world fisheries. Nat. Clim. Chang. 1, 449–456.

Sweeney, C., Gnanadesikan, A., Griffies, S.M., Harrison, M.J.,

Rosati, A.J., Samuels, B.L., 2005. Impacts of shortwave penetration

depth on large-scale ocean circulation and heat transport. J. Phys.

Oceanogr. 35, 1103–1119.

Takahashi, T., Sutherland, S.C., Wanninkhof, R., Sweeney, C.,

Feely, R.A., Chipman, D.W., Hales, B., Friederich, G., Chavez, F.,

Sabine, C., Watson, A., Bakker, D.C.E., Schuster, U., Metzl, N.,

Inoue, H.Y., Ishii, M., Midorikawa, T., Nojiri, Y., Koertzinger, A.,

Steinhoff, T., Hoppema, M., Olafsson, J., Arnarson, T.S.,

Tilbrook, B., Johannessen, T., Olsen, A., Bellerby, R., Wong, C.S.,

Delille, B., Bates, N.R., de Baar, H.J.W., 2009. Climatological mean

and decadal change in surface ocean pCO2, and net sea-air CO2 flux

over the global oceans. Deep Sea Res. Part II 56, 554–577.

Tanhua, et al., 2013. The carbon cycle and the ocean. In: Siedler, G.,

Church, J., Gould, J., Griffies, S. (Eds.), Ocean Circulation and

Climate. second ed. Academic Press (Elsevier), Waltham, MA, USA.

Tans, P., Keeling, R., 2012. Trends in atmospheric carbon dioxide, full

Mauna Loa CO2 record. http://www.esrl.noaa.gov/gmd/ccgg/trends/,

Dr. Pieter Tans, NOAA/ESRL (www.esrl.noaa.gov/gmd/ccgg/

trends/) and Dr. Ralph Keeling, Scripps Institution of Oceanography

(scrippsco2.ucsd.edu/) (Accessed Dec. 2012).

Thomas, M.K., Kremer, C.T., Klausmeier, C.A., Litchman, E., 2012. A

global pattern of thermal adaptation in marine phytoplankton. Science

338, 1085–1088.

Toole, D.A., Siegel, D.A., 2004. Light-driven cycling of dimethylsulfide

(DMS) in the Sargasso Sea: closing the loop. Geophys. Res. Lett.

31, L09308. http://dx.doi.org/10.1029/2004GL019581.

Toole, D.A., Siegel, D.A., Doney, S.C., 2008. A light-driven, one-

dimensional dimethylsulfide biogeochemical cycling model for the

Sargasso Sea. J. Geophys. Res. 113, G02009. http://dx.doi.org/

10.1029/2007JG000426.

VanMooy,B.A.S.,Fredricks,H.F., Pedler, B.E., Dyhrman, S.T., Karl,D.M.,

Koblızek, M., Lomas, M.W., Mincer, T.J., Moore, L.R., Moutin, T.,

Rappe, M.S., Webb, E.A., 2009. Phytoplankton in the ocean use non-

phosphorus lipids in response tophosphorus scarcity.Nature 458, 69–72.

Vogt, M., Liss, P.S., 2009. Dimethylsulfide and climate. In: Le Quere, C.,

Saltzman, E.S. (Eds.), Surface Ocean—Lower Atmosphere Processes.

Geophysical Monograph Series, vol. 187. American Geophysical

Union, Washington, DC, pp. 197–232. http://dx.doi.org/10.1029/

2008GM000790.

Volk, T., Hoffert, M.I., 1985. Ocean carbon pumps: analysis of relative

strengths and efficiencies in ocean-driven atmospheric CO2 changes.

PART VI The Changing Ocean842

In: The Carbon Cycle and Atmospheric CO2: Natural Variations

Archean to Present. GeophysicalMonograph Series, vol. 32. American

Geophysical Union, Washington, DC, USA, pp. 99–110.

Ward, J.R., Lafferty, K.D., 2004. The elusive baseline of marine disease:

are diseases in ocean ecosystems increasing? PLoS Biol. 2, e120.

Wassmann, P., Duarte, C.M., Agusti, S., Sejr, M.K., 2011. Footprints of

climate change in the Arctic marine ecosystem. Glob. Chang. Biol 17,

1235–1249.

Waycott,M.,Duarte, C.M., Carruthers, T.J.B., Orth, R.J., Dennison,W.C.,

Olyarnik, S., Calladine, A., Fourqurean, J.W., Heck Jr., K.L.,

Hughes, A.R., Kendrick, G.A., Kenworthy, W.J., Short, F.T.,

Williams, S.L., 2009. Accelerating loss of seagrasses across the globe

threatens coastal ecosystems. Proc. Natl. Acad. Sci. U.S.A. 106,

12377–12381.

Weil, E., Croquer, A., Urreiztieta, I., 2009. Temporal variability and con-

sequences of coral diseases and bleaching in La Parguera, Puerto Rico

from 2003–2007. Caribb. J. Sci. 45, 221–246.

White, A.E., Bjorkman, K., Grabowski, E., Letelier, R.M., Poulos, S.,

Watkins, B., Karl, D.M., 2010. An open ocean trial of controlled

upwelling using wave pump technology. J. Atmos. Oceanic Technol.

27, 385–396.

Whitney, F.A., Freeland, H.J., Robert, M., 2007. Persistently declining

oxygen levels in the interior waters of the eastern subarctic Pacific.

Prog. Oceanogr. 75, 179–199.

Williams, R.G., Follows, M.J., 2011. Ocean Dynamics and the Carbon

Cycle. Cambridge University Press, Cambridge, UK, 404 pp.

Wilson, S., Fischetti, T., 2010. Coastline population trends in the

United States: 1960 to 2008. In: Current Population Reports: Estimates

and projections, P25-1139. US Census Bureau, Washington DC,

27 pp.

Wolf, S.G., Snyder, M.A., Sydeman, W.J., Doak, D.F., 2010. Predicting

population consequences of ocean climate change for an ecosystem

sentinel, the seabird Cassin’s auklet. Glob. Chang. Biol 16,

1923–1935.