Plant, Cell and Environment (1999) 22, 741–755
ORIGINAL ARTICLE
OA
220
EN
Oceanic sinks for atmospheric CO2
J. A. RAVEN 1 & P. G. FALKOWSKI 2
Department of Biological Sciences, University of Dundee, Dundee DD1 4HN, UK and 2Institute of Marine and Coastal Sciences,
Rutgers University, New Brunswick, NJ 08903, USA
1
ABSTRACT
There is approximately 50 times more inorganic carbon
in the global ocean than in the atmosphere. On time
scales of decades to millions of years, the interaction
between these two geophysical fluids determines atmospheric CO2 levels. During glacial periods, for example,
the ocean serves as the major sink for atmospheric CO2,
while during glacial–interglacial transitions, it is a source
of CO2 to the atmosphere. The mechanisms responsible
for determining the sign of the net exchange of CO2
between the ocean and the atmosphere remain unresolved. There is evidence that during glacial periods,
phytoplankton primary productivity increased, leading
to an enhanced sedimentation of particulate organic carbon into the ocean interior. The stimulation of primary
production in glacial episodes can be correlated with
increased inputs of nutrients limiting productivity, especially aeolian iron. Iron directly enhances primary production in high nutrient (nitrate and phosphate) regions
of the ocean, of which the Southern Ocean is the most
important. This trace element can also enhance nitrogen
fixation, and thereby indirectly stimulate primary production throughout the low nutrient regions of the central ocean basins. While the export flux of organic
carbon to the ocean interior was enhanced during glacial
periods, this process does not fully account for the
sequestration of atmospheric CO2. Heterotrophic oxidation of the newly formed organic carbon, forming weak
acids, would have hydrolyzed CaCO3 in the sediments,
increasing thereby oceanic alkalinity which, in turn,
would have promoted the drawdown of atmospheric
CO2. This latter mechanism is consistent with the stable
carbon isotope pattern derived from air trapped in ice
cores. The oceans have also played a major role as a sink
for up to 30% of the anthropogenic CO2 produced during
the industrial revolution. In large part this is due to CO2
solution in the surface ocean; however, some, poorly
quantified fraction is a result of increased new production due to anthropogenic inputs of combined N, P and
Fe. Based on ‘circulation as usual’, models predict that
future anthropogenic CO2 inputs to the atmosphere will,
in part, continue to be sequestered in the ocean. Human
intervention (large-scale Fe fertilization; direct CO2
burial in the deep ocean) could increase carbon sequestration in the oceans, but could also result in unpre-
dicted environmental perturbations. Changes in the
oceanic thermohaline circulation as a result of global
climate change would greatly alter the predictions of C
sequestration that are possible on a ‘circulation as usual’
basis.
Key-words: CO2; carbon concentrating mechanisms (CCMs);
Fe; N; ocean circulation; P; solubility; temperature.
INTRODUCTION
The oceans cover 70% of the Earth’s surface and contain
about 50 times more soluble inorganic carbon than the
atmosphere. The fluxes of CO2 between these two geophysical fluids are enormous, averaging approximately
100 Pg C per annum. Since the Industrial Revolution,
there has been a net flux of CO2 from the atmosphere to
the oceans, which presently amounts to about 2 Pg per
annum. Up to an additional 1·5 Pg C may be sequestered
by terrestrial ecosystems. On time scales of millions of
years, the net flux of CO2 between the oceans and atmospheres is driven by very small changes in both inorganic
and biological processes. On time scales of years to
decades, the net flux between the oceans and atmosphere
is primarily driven by physico-chemical processes related
to solubilization of CO2 in the oceans, while on time
scales of centuries to millennia biological responses play
a key role (see review by Sarmiento & Bender 1994).
Geochemical models clearly suggest that removal of
atmospheric CO2 by a terrestrial sink (e.g. enhanced net
primary production) will result in outgassing of CO2
from the oceans. We strive here to explain the critical factors that determine the net fluxes of CO2 between the
oceans and atmosphere, with a goal of explaining key feedbacks in the global carbon cycle. We begin by considering
the global carbon cycle in terms of the chemical components and the biogeochemical mechanisms involved, the
pool sizes of the components and the magnitude of the
fluxes among them.
THE ROLE OF THE OCEANS IN THE NATURAL
GLOBAL CARBON CYCLE
Correspondence: J. A. Raven
On geological time scales (greater than hundreds or thousands of years), inorganic carbon is sequestered in the
oceans as either Ca or Mg mineral salts. The geochemical
reaction schemes responsible for the formation of the Ca
and Mg mineral salts were first described by Harold Urey
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J. A. Raven and P. G. Falkowski
(Urey 1952; see also Berner, Lasaga & Garrels 1983;
Berner & Berner 1996). These so-called ‘Urey reactions’
are summarized thus:
CO2 + CaSiO3
CO2 + MgSiO3
weathering
→
←
metamorphism
weathering
→
←
metamorphism
CaCO3 + SiO2
(1)
MgCO3 + SiO2
(2)
It should be noted that CaSiO3 and MgSiO3 represent any
generic Ca or Mg silicate produced at high temperatures in
the Earth’s crust, that SiO2 represents any sedimented silicate which has not been subjected to high temperatures,
and that MgCO3 is the Mg component of the mixed Mg
and Ca carbonate, dolomite (Berner et al. 1983). In Eqns 1
and 2 ‘weathering’ occurs on the land surface (noting that
continental crust has existed for as long as we have fossil
evidence of cyanobacteria, i.e. 3·45 billion years: Raven
1998a) and yields bicarbonates rather than carbonates. The
precipitation of CaCO3 and MgCO3 (and SiO2) occurs
only in the oceans. Thus, the terrestrial part of the weathering reactions in Eqns 1 and 2 may be represented as:
weathering
→
Ca(HCO3)2 + SiO2 (3)
2 CO2 + H2O + CaSiO3
weathering
Mg(HCO3)2 + SiO2 (4)
2 CO2 + H2O + MgSiO3 →
While these reactions are inorganic, they are influenced by
biological processes. With the emergence of terrestrial
vegetation as cyanobacteria and microalgae 1·2 billion
years ago, and especially as higher plants (embryophytes)
420 million years ago (Raven 1998), the weathering
reactions were accelerated. Terrestrial plants fix atmospheric CO2, and transfer a fraction of the fixed organic
carbon to the soil surface, where it is respired. This terrestrial CO2 pump enriches groundwater, and hence rivers,
with inorganic carbon, with little possibility of diffusive
efflux to the atmosphere (Berner et al. 1983; Berner 1992,
1994, 1997, Retallack 1997; Algeo & Schleckler 1998;
Berner 1998; Elick, Driese & Mora 1998). The flux of
inorganic carbon and its associated divalent cations
enriches the oceans with these inorganic molecules on time
scales of millennia.
Following the transport of Ca(HCO3)2, Mg(HCO3)2 and
SiO2 (as (Si(OH)4), from fluvial sources to the ocean, precipitation reactions become dominant. The marine precipitation processes are represented as:
marine precipitation
→
CaCO3 + CO2 + H2O (5)
Ca(HCO3)2
marine precipitation
→
MgCO3 + CO2 + H2O (6)
Mg(HCO3)2
The sum of Eqns 3 and 5 is the weathering component of
Eqn 1, while the sum of Eqns 4 and 6 is the weathering
component of Eqn 2. In the contemporary oceans there is
very little precipitation of dolomite; and essentially all of
the precipitation of CaCO3 (as calcite) and SiO2 is biologically mediated (Berner et al. 1983; Berner & Berner 1996).
Another component of Urey reactions is the sequestration of CO2 on land by the dissolution of carbonates that
had precipitated previously in the ocean according to
Eqns 5 and 6, but escaped metamorphism and subsequently were uplifted and exposed. This carbonate dissolution pathway effectively is the reverse of Eqns 5 and 6,
thus:
CaCO3 + CO2 + H2O
MgCO3 + CO2 + H2O
dissolution
→
dissolution
→
Ca(HCO3)2
(7)
Mg(HCO3)2
(8)
As with the silicate weathering in Eqns 3 and 4, these
reactions are enhanced by biological regeneration in soils
(Berner et al. 1983; Berner 1997, 1998; Retallack 1997).
The estimated fluxes of CO2 associated with Urey
reactions were calculated by Berner et al. (1983). The
uptake of CO2 from the atmosphere related to carbonate
dissolution on land (Eqns 7 and 8) is 0·14 Pg C per year,
and that associated with silicate weathering (Eqns 3 and
4) is 0·14 Pg C per year, yielding a total of 0·28 Pg C per
year. Assuming a steady-state, CO2 input resulting from
calcite precipitation in the ocean (Eqns 5 and 6) amounts
to 0·209 Pg C per year. Vulcanism and other metamorphic fluxes (the ‘metamorphic’ direction of Eqns 1 and
2) produce 0·071 Pg C per year. Hence, the total of
0·28 Pg C per year.
These global fluxes, driven by geochemical processes
are relatively small compared with biological fluxes of
carbon. Terrestrial net primary production (NPP) is
presently estimated at approximately 56 Pg C per year,
while marine NPP is approximately 45 Pg C per year
(Field et al. 1998). However, small changes in the Urey
reactions are very significant as long-term (millions to
billions of years) determinants of atmospheric CO2
(Berner et al. 1983; Berner 1994, 1997; Retallack 1997).
Among the more important processes altering the rates
of CO2 input to the atmosphere by metamorphism
(Eqns 1 and 2) is the extent of tectonic plate subduction.
More subduction leads to a greater magmatic production
of CO2 by metamorphism (Berner et al. 1983). CO2
removal by the weathering processes described in
Eqns 3, 4, 7, and 8 is also influenced by plate tectonics in
that tectonic processes determine the continental area
exposed to weathering (Berner et al. 1983).
Much of the rest of this paper is concerned with the factors controlling primary productivity in the long term
sequestration of carbon in the oceans. Our approach uses
palaeo-ecological data on the role of oceans and their biota
in previous episodes of variation in atmospheric CO2,
especially the last glaciation, as well as mechanistic models of what may happen as CO2 increases. We also consider
interventionist procedures such as iron fertilization and
direct transfer of CO2 to the deep ocean.
© 1999 Blackwell Science Ltd, Plant, Cell and Environment, 22, 741–755
Oceanic sinks for atmospheric CO2
CONSTRAINTS ON PRIMARY PRODUCTIVITY
AND ITS RECYCLING TO CO 2 IN TODAY’S
OCEAN AND ITS IMPLICATIONS FOR
ATMOSPHERIC CO 2
Supply of inorganic and organic C from
terrestrial biota
The ocean is a net recipient of both organic and inorganic
carbon from terrestrial systems. Respiration in soils
inevitably leads to the production of CO2, which equilibrates between the solution and gas phases. While diffusion of some CO2 gas to the atmosphere occurs, some of
the dissolved CO2 becomes hydrated and promotes the
weathering reactions depicted in Eqns 3, 4, 7 and 8. The
hydrated CO2 (i.e. H2CO3 + HCO3–+ CO32–), together
with some of the free dissolved CO2 is ultimately transported to bodies of freshwater where biologically mediated
oxidation of allochthonously produced (i.e. imported)
organic C generates more CO2. In open freshwater aquatic
ecosystems, CO2 can become supersaturated and, despite
photosynthetic consumption, can outgas to the atmosphere
(Cole et al. 1994). The net effect of these processes is a
flux of CO2 from freshwaters to the atmosphere amounting
to approximately 0·14 Pg C (Cole et al. 1994).
Essentially all of the HCO3– (0·28 Pg C) generated in
the reactions defined in Eqns 3, 4, 7 and 8 is transported
to the ocean in rivers. The flux of organic carbon
amounts to approximately 0·4 Pg C (see Meybeck 1993;
Watson & Liss 1998). The terrestrially derived organic
carbon is often identified by the presence of lignin
(which is not significantly produced by marine photoautotrophs). In the oceans, lignin, which is relatively
resistant to decomposition, is found only in coastal environments, especially in sediments of continental margins
(Premuzic et al. 1982). Hence, the coastal oceans are the
primary sink for terrestrially derived organic matter
exported from the continents. The total flux of inorganic
plus organic carbon from terrestrial to marine ecosystems is approximately 0·7 Pg C per annum. The inorganic component helps to subsidize marine primary
productivity, while the organic carbon component is of a
similar magnitude to the long-term storage of organic
carbon in oceans (Beran 1995).
Ocean circulation and distribution of CO2 and
other solutes
On time scales of centuries to millennia, oceanic inorganic
C controls the atmospheric level of CO2, not vice versa.
The equilibrium distribution of CO2 between the ocean and
the atmosphere is critically dependent upon the temperature, alkalinity and salinity of surface waters (Broecker
1985; Takahashi 1989; Bigg 1996; Sarmiento & Quéré
1996; Broecker 1997). Mean oceanic salinity and alkalinity vary slightly with freshwater sequestration in ice sheets
over millennial time-scales, while mean sea surface temperature varies over a shorter time scale. Higher temperatures and salinity mean lower CO2 solubility.
© 1999 Blackwell Science Ltd, Plant, Cell and Environment, 22, 741–755
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There are variations in temperature, alkalinity and salinity driven by solar radiation via the hydrological cycle and
the warming of low-latitude oceans. In the absence of any
biological activity in the oceans, water which cools and
then downwells at high latitudes has high CO2 concentrations (not necessarily in full equilibrium with the atmosphere: Smith 1985; Watson, Upstill-Goddard & Liss
1991). This situation occurs only in the North Atlantic, off
the coast of Greenland and Labrador, and in the Southern
Ocean, but not in the North Pacific. Upon upwelling at low
latitudes the cold, CO2-enriched, water from the ocean
interior warms and becomes supersaturated with respect to
CO2. Prior to the Industrial Revolution, the net flux of carbon from high latitudes to low latitudes in the ocean
amounted to some 1 Pg C per year (see Watson & Liss
1998). This flux was balanced by the transport of 1 Pg C
from low to high latitudes in the atmosphere (Fig. 1). The
oceanic absorption of CO2 at high latitudes and subsequent
subduction is called the ‘solubility pump’, and is primarily
maintained by cold temperatures in the ocean interior.
Consequently, the concentration of inorganic carbon below
the upper 500 m of the ocean is significantly higher than
the air–sea equilibrium values (Fig. 2).
Figure 1. The transport of carbon between low and high latitudes
prior to (a), and subsequent to (b), the Industrial Revolution. Prior
to the Industrial Revolution, the atmosphere and ocean were in
steady-state on the time scale of decades, and approximately 1 Pg
of carbon (as CO2) was injected each year into the atmosphere
from upwelling systems at low latitudes, while an equivalent
amount was returned to the oceans in the formation of cold, dense
water at high latitudes (Fig. 1a). In the contemporary ocean, the
imbalance in the carbon cycle resulting from anthropogenic
activities leads effectively to a decrease in the uptake of CO2 by the
oceans, with only 0·7 Pg of carbon per year transferred to the ocean
from the atmosphere at high latitudes whilst 1 Pg of carbon is
injected into the atmosphere by low-latitude upwelling systems
(Fig. 1b). See Watson & Liss (1998).
744
J. A. Raven and P. G. Falkowski
Figure 2. Vertical profiles of total dissolved inorganic carbon
(TIC) in the ocean. Curve A corresponds to a theoretical profile that
would have been obtained prior to the Industrial Revolution with an
atmospheric CO2 concentration of 280 µmol mol–1. The curve is
derived from the solubility coefficients for CO2 in seawater, using a
typical thermal and salinity profile from the central Pacific Ocean,
and assumes that when surface water cools and sinks to become
deep water it has equilibrated with atmospheric CO2. As such, the
calculated profile of TIC reflects the ‘solubility pump’ and assumes
that the ‘biological pump’ is nil. Curve B corresponds to the same
calculated solubility profile of TIC, but in the year 1995, with an
atmospheric CO2 concentration of 360 µmol mol–1. The difference
between these two curves is the integrated oceanic uptake of CO2
from anthropogenic emissions since beginning of the Industrial
Revolution, with the assumption that biological processes have
been in steady-state (and hence have not materially affected the net
influx of CO2). Curve C is a representative profile of measured TIC
from the central Pacific Ocean. The difference between curve C
and B is the contribution of biological processes to the uptake of
CO2 in the steady-state (i.e. the contribution of the ‘biological
pump’ to the TIC pool. (Data courtesy of Doug Wallace and the
World Ocean Circulation Experiment).
Oceanic photosynthesis further modifies this situation. A
fraction of the carbon fixed by phytoplankton in the upper
ocean is exported (sinks) to the ocean interior, where it is
oxidized by microbes, regenerating CO2 in the deep ocean.
This ‘biological pump’ (Volk & Hoffert 1985) produces a
vertical-inorganic carbon profile that is further enriched at
depth (Fig. 2). A similar profile is found for other biologically active solutes in oceans (e.g. NO3–, NH4+, HPO42–,
Si(OH)4, and Fe). Unlike other nutrients, however, the
fractional difference in total inorganic carbon concentrations between surface and deeper water is much smaller
than for NO3–, NH4+ and HPO42–, providing prima facie
evidence that inorganic carbon does not limit primary productivity in the oceans.
Physical processes that influence vertical fluxes of nutrients are the critical determinants of primary production in
the ocean. Solar heating of surface ocean waters isolates
the upper mixed layer from the cold, deep ocean interior.
The thermal gradient constrains phytoplankton in a wellilluminated (euphotic) zone, but retards mixing with
deeper, nutrient-rich waters. This situation permanently
characterizes tropical and subtropical seas. Seasonal
heating and cooling in temperate and boreal regions,
coastal upwelling driven by winds, and the dissipation of
kinetic energy by turbulent mixing (e.g. storms, eddies,
and high frequency internal waves), facilitates the injection of nutrients into the euphotic zone from deeper
waters. The importation of ‘new’ nutrients into the
euphotic zone permits transient blooms of phytoplankton.
In the steady-state these fluxes are balanced by the vertical ‘export’ flux of organic material to the ocean interior.
This export flux fuels biological processes in the deep
ocean. Alternatively, phytoplankton can be consumed
within the euphotic zone, leading to a ‘recycling’ of nutrients. About 70% of global marine primary productivity is
‘recycled’ and some 30% is ‘new’ (Falkowski & Raven
1997; Field et al. 1998). While distribution of new productivity is highly variable spatially and temporally, it is
important to emphasize that neither ‘new’ nor recycled
production affects the global net exchange of CO2
between the atmosphere and ocean.
The average elemental ratio of inorganic C : N : P regenerated by the oxidation of organic matter in the ocean is
highly constrained. This so-called ‘Redfield ratio’ of
106C : 16N : 1P (by atoms), is unique to marine ecosystems and reflects the annually and spatially averaged,
proximate, highly conserved, chemical composition of
phytoplankton (Redfield 1958; Copin-Montegut & CopinMontegut 1983). The Redfield ratio is useful in calculating
the organic carbon potentially produced for a given amount
of fixed inorganic nitrogen and phosphate, regardless of
whether the production is regenerated or ‘new’.
Throughout most of the central oceans, the ratio of
fixed inorganic nitrogen (in the form of NO3–) to P (as
HPO42–) averages 14·7 rather than 16 (Falkowski 1997a;
Gruber & Sarmiento 1997). This small difference is a
consequence of losses of combined nitrogen due to denitrification relative to nitrogen inputs via N2 fixation or
riverine fluxes (Codispoti 1995). Carbon and phosphorus, and to a much lesser extent, nitrogen, can be lost by
sedimentation (Berner 1980). Prior to the Industrial
Revolution, the quantity of organic carbon buried in
ocean sediments amounted to some 0·2 Pg C per year
(Fig. 1 of Watson & Liss 1998; cf. Chin, Orelland &
Verdugo 1998; Wells 1998). This mass balance estimate
implies that less organic carbon is sedimented in the central oceans basins than is delivered by rivers. Virtually
all of the oceanic sedimentation of organic carbon occurs
in coastal waters (see Hartnett et al. 1998; Reimers
1998), and a significant fraction of this organic carbon is
derived from terrestrial primary productivity. The accumulation of organic carbon in ocean margin sediments
represents very small imbalances between organic carbon production and respiration. Almost no terrestrial or
aquatic NPP escapes oxidation to CO2 either biologically
© 1999 Blackwell Science Ltd, Plant, Cell and Environment, 22, 741–755
Oceanic sinks for atmospheric CO2
through heterotrophic respiration or, for some terrestrial
habitats, by combustion from fires. The average turnover
time of carbon in terrestrial ecosystems is on the order of
two decades, whilst in aquatic ecosystems the average
turnover time is on the order of a week (Beran 1995;
Falkowski, Barber & Smetacek 1998).
To summarize, it is generally assumed that prior to the
Industrial Revolution, the global carbon cycle was
‘steady state’ during the relatively brief (on geological
time scales) period from 10 000 to 250 years before the
present. There was a net flux of CO2 from the oceans to
the atmosphere amounting to some 0·5 Pg C per year
during this period. This flux was balanced by an equal
flux of CO2 into organic matter in terrestrial biota, and
thence back to the oceans in rivers. Approximately 1 Pg
C (as inorganic carbon) flowed from high latitudes to
low latitudes in the ocean. An equal flux of CO2 was conveyed through the atmosphere from warm, low-latitude
regions to colder high-latitudes (Fig. 1). Approximately
0·2 Pg C as CO2 supplied to the atmosphere from vulcanism was incorporated into terrestrial biota and thence,
via rivers, to marine sediments (and thus back to the C
source for CO2 emitted in vulcanism): Watson & Liss
(1998) (Table 1). The geochemical assumption of a
steady-state carbon cycle for at least 10 000 years prior
to the Industrial Revolution is predicated on the principle
that a deviation from steady-state requires that at least
one of three conditions must be met: (1) nutrients limiting primary production must be added to the ocean from
external sources; (2) limiting nutrients in the euphotic
zone that are unused on an annual basis must be consumed, and/or (3) the chemical composition of phytoplankton must change.
‘Bottom-up’ versus ‘top-down’ control of marine
primary productivity
The standing stock of photosynthetic biomass in the world
oceans is extremely small relative to terrestrial ecosystems. The estimated oceanic photosynthetic biomass of
approximately 1 Pg C amounts to approximately 0·2% of
the total photosynthetic biomass on Earth (Falkowski &
Raven 1997; Field et al. 1998). The rapid turnover of this
biomass is related to both the rate of production and the
rate of removal. The balance between these two processes
as a function of time can be described by the simple differential equation:
dN/dt = N(µ – m)
where N is the ensemble of photosynthetic biomass, µ is
the specific growth rate, and m is the specific mortality.
Both µ and m have units of (time)–1.
Any factor that limits the intrinsic physiological rate of
growth, expressed as µ, is considered ‘bottom-up’ control.
Such a factor may be a limiting resource (light, C, N, P, Fe,
Si, vitamins, etc.) or a physical condition (thermal energy,
osmolality). Such restrictions on the rate of biomass
increase are termed ‘stress’ by Grime (1979). The history
of bottom-up control can be traced to von Liebig (1840),
with what is now termed ‘extent’ limitation (as in a batch
culture). The concept of ‘rate’ limitation (as in a chemostat
or turbidostat) is generally attributed to Blackman (1905)
and Nathansohn (1908) (de Baar 1994; Falkowski 1994).
Primary production in the oceans is generally limited
both in extent and rate by nutrient fluxes (i.e. ‘bottom-up’
control). Throughout most of the upper ocean, the concentrations of total fixed inorganic nitrogen and phosphorus
Flux
Pg C year–1
Volcanic/tectonic to atmosphere
From atmosphere in terrestrial gross primary productivity
From land biota to atmosphere by respiration and fire
From soil to atmosphere by respiration
From soil in rivers
From atmosphere to warm surface ocean
From warm surface ocean to atmosphere
From warm surface ocean to biota in gross primary productivity
From biota to warm surface ocean in respiration
From atmosphere to cold surface ocean
From cold surface ocean to atmosphere
From cold surface ocean to biota in gross primary productivity
From biota to cold surface ocean in respiration
From warm surface ocean to cold surface ocean as inorganic C
From cold surface ocean to intermediate and deep waters as inorganic C
From warm surface ocean biota to intermediate and deep waters
as organic C and CaCO3
From intermediate and deep waters to warm surface ocean as inorganic C
From cold surface ocean biota to intermediate and deep waters
as organic C and CaCO3
From intermediate and deep waters to cold surface ocean as inorganic C
From biota in intermediate and deep waters to sediments as organic
C and CaCO3
0·2
100·7
50
50
0·7
70
71·5
20
18
20
19
10
9
19·2
50·2
2
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1
31
0·2
Table 1. Carbon fluxes in the pre-industrial
(up to approximately 1750 AD) part of the
present interglacial. Adapted from Fig. 1 of
Watson & Liss (1998). Note that, to ensure
closure of cycles, the values differ slightly
from those in the text
746
J. A. Raven and P. G. Falkowski
are extremely low (< 1 mmol m–3). Vertical profiles of
these two nutrients clearly reveal that fixed inorganic nitrogen is almost always depleted more rapidly than phosphate
in the euphotic zone (Fanning 1992), and indeed, to a first
approximation, fixed nitrogen limits phytoplankton
biomass and growth throughout most of the world oceans
(Falkowski et al. 1998; Downing 1997; notwithstanding).
The paucity of nitrogen is curious. In most temperate
and tropical lakes, when fixed inorganic nitrogen concentrations become low, nitrogen fixation is stimulated.
Consequently, phosphorus, which, unlike nitrogen, has no
atmospheric source, almost always limits primary production in lakes (Hutchinson 1957). In the ocean, however,
diazotrophic organisms are relatively sparse and confined
to very few taxa. The lack of both numbers and diversity of
nitrogen fixers in marine ecosystems has led to the notion
that nitrogen limitation is itself limited by some other factor, the most probable being iron (Falkowski 1997a).
A major source of iron for the oceans is aeolian dust
(Duce 1986). Ferric iron is virtually insoluble in seawater,
and precipitates as hydrated oxides or phosphates. The
concentration of iron throughout most of the euphotic zone
in the world’s oceans is approximately 1 mmol m–3 or less.
Terrestrially derived iron, primarily from deserts, is delivered by prevailing winds, and provides the essential transition metal for phytoplankton growth (Blank, Leinen &
Prospero 1985; Prospero & Nees 1985). Nitrogen fixers
have an extremely high iron requirement, and in oceanic
areas far removed from an aeolian source, iron fluxes cannot sustain high rates of fixation. The iron limitation of
nitrogen fixation is particularly acute in the southern hemisphere, especially in the Pacific.
In three regions of the world oceans, namely the eastern
equatorial Pacific, the subarctic Pacific, and the Southern
(Antarctic) Oceans, iron limitation is so severe that photosynthetic electron transport is limited by the availability of
the metal (Kolber et al. 1994; Behrenfield et al. 1996; P.
Boyd, personal communication). In these three regions,
iron limitation actually prevents the utilization of fixed
nitrogen and phosphorus in the euphotic zone. Hence, the
concentration of NO3– and HPO2–4 in these regions can be
relatively high, approaching 30 to 50 mmol m–3 in some
areas. Stimulation of primary production by the addition of
iron to these so-called ‘high nutrient (N, P)-low chlorophyll’ regions of the ocean can significantly influence
atmospheric CO2 influx (Martin, Gordon & Fitzwater
1990;Falkowski 1994; de Baar et al. 1995; Raven 1995;
Coale et al. 1996; Cooper, Watson & Nightingale 1996;
Pakulski et al. 1996; Falkowski 1997a; cf. Hart 1934;
Harvey 1937).
Specific taxa are sometimes limited by specific elements. Diatoms are perhaps the most important taxon
mediating the export flux of carbon in the oceans. This
group of organisms uses silica to make opaline cell walls.
Silica is supplied to the oceans from fluvial sources and the
element often limits diatom production in coastal waters
(Dugdale & Wilkerson 1998; Smetacek 1998). Vitamins
have been inferred to limit dinoflagellates and other groups
(Provasoli & Carlucci 1974). While these compounds and
other specific elements, such as Cu, Co, Mo and Zn, are
sometimes implicated as factors selecting specific phytoplankton taxa, they do not limit overall NPP in the oceans
(Morel et al. 1994).
The mortality of photoautotrophs involves grazers and
parasites, namely processes dependent on other trophic
levels. Regulation of biomass via such processes is termed
‘top-down’ control. Mortality is essential to the recycling
of nutrients, which permits the growth of the remaining
photoautotrophic biomass; however, it cannot alter the
maximum biomass supported by nutrient supply. Hence,
while ‘top-down’ control plays an extremely important role
in determining the structure of marine food webs and nutrient recycling efficiency (Banse 1992; Azam 1998), from a
geochemical perspective, it plays a relatively unimportant
role in regulating the net exchange of CO2 between the
oceans and atmosphere.
Influence of increased atmospheric CO2 on
marine primary productivity via increases in
surface ocean inorganic carbon
With an atmospheric CO2 concentration of 365 µmol mol–1
(a typical value over the last decade), the equilibrium value
for total inorganic carbon at the ocean surface is approximately 2 mol m–3 at 18 °C (the mean global sea surface
temperature). Approximately 95% of the inorganic carbon
in the ocean is in the form of bicarbonate anion; the equilibrium concentration of CO2 is only approximately
10 mmol m–3. This concentration of CO2 is about an order
of magnitude lower than that required to saturate
RuBisCO, and CO2 is frequently below air-equilibrium in
regions with high levels of other (noncarbon) nutrients
(Codispoti et al. 1982; Codispoti, Friedrich & Hood 1986;
Watson et al. 1991a,b). Hence, in the absence of either an
alternative carbon fixing pathway or a CO2 concentrating
mechanism, marine phytoplankton (and macrophyte) photosynthesis at light saturation in air-equilibrated seawater
would be limited by inorganic carbon (Raven 1970).
Over the past two decades, it has become clear that many
marine photoautotrophs can concentrate inorganic carbon
via ‘carbon concentrating mechanisms’ (CCMs) (Paasche
1964; Thomas & Tregunna 1968; Pruder & Bolton 1979;
Miller, Turpin & Canvin 1984; Badour & Irvine 1990;
Raven 1991; Raven & Johnston 1991; Raven, Johnston &
Turpin 1993; Riebesell, Wolf-Gladrow & Smetacek 1993;
Kübler & Raven 1994; Morel et al. 1994; Thom 1995;
Beer & Koch 1996; Berry et al. 1996; Kübler & Raven
1996; Paasche et al. 1996; Uusitalo 1996; Giordano &
Bowes 1997; Hassidim et al. 1997; Hein & Sand-Jensen
1997; Raven 1997; Tortell, Reinfelder & Morel 1997;
Beardall, Johnston & Raven 1998a; Beardall, Beer &
Raven 1998b;Sültemeyer et al. 1998). The presence of
CCMs can facilitate the influx of either CO2 or HCO3–
(Raven 1997; Wolf-Gladrow & Riebesell 1997; Kaplan
et al. 1998; cf. Kaneko & Table 1997), and differential CO2
and HCO3– uptake can lead to disequilibrium between CO2
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Oceanic sinks for atmospheric CO2
and HCO3– in the surface ocean during phytoplankton
blooms. There is a small efflux (leak) of inorganic carbon,
as well as an influx of inorganic carbon during the operation of CCMs (Raven 1997). While the magnitude of CO2
efflux at high photon flux densities can be relatively large
under contrived laboratory conditions (Sukenik et al. 1997;
Tchernov et al. 1997), this flux is generally much too small
to be either biologically or geochemically significant
(Falkowski 1997b).
A large number of experiments have been conducted to
examine potential limitation of photosynthesis by inorganic carbon in marine photoautotrophs. Most of the currently available data concern short-term (minutes – tens of
minutes) measurements of photosynthetic rates as a function of inorganic carbon concentration for organisms
grown at or near air-equilibrium CO2 levels in seawater
enriched with other nutrients. Many of these data may not
reflect organisms grown at air-equilibrium, since growth
procedures frequently allow CO2 (and other components of
the inorganic carbon system) to be depleted. Despite the
difficulty in constraining all potential experimental variables, the outcome of these measurements is that the lightsaturated or light-limited rate of photosynthesis is not
limited by inorganic carbon in equilibrium with the atmosphere for all marine cyanobacteria, most eukaryotic
microalgae and many eukaryotic macroalgae (Muñoz &
Merrett 1989; Riebesell et al. 1993; Kübler & Raven 1994;
Beer & Koch 1996; Kübler & Raven 1996; Hein & SandJensen 1997; Raven 1997; Tortell, Reinfelder & Morel
1997; Beardall et al. 1998b). Virtually all marine
embryophytes (i.e. seagrasses) are limited by inorganic
carbon (Beer & Koch 1996; Raven 1997; Zimmerman
et al. 1997; Beardall et al. 1998b). These results strongly
suggest that changes in atmospheric CO2 will not (and historically, have not) directly affect overall photosynthetic
rates in marine planktonic ecosystems.
For growth there are many fewer data, and there are reasons to believe that growth will not have a higher inorganic
C affinity than does short-term photosynthesis (Raven,
Johnston & Turpin 1993; Falkowski & Raven 1997). This
prediction is largely borne out by the available data (Thom
1995; Paasche et al. 1996; Raven 1997). More data are
needed on growth as a function of inorganic C under light,
nitrogen, phosphorus or iron-limited conditions which are
commonly found in nature (Kübler & Raven 1994, 1996).
The inorganic carbon system in the oceans is the primary
pH buffer. This buffer is dependent upon the availability of
alkaline earth cations, especially Ca2+. The precipitation of
carbonate is an ancient pathway that has provided most of
the 60 000 000 Pg of C as CaCO3 in the lithosphere. In the
open ocean, primary producers, such as coccolithophorids
and symbiotic foraminifera, are a significant source of
CaCO3 in the form of calcite. In coastal waters, CaCO3precipitating primary producers include symbiotic corals
and foramenifera, as well as red, green and brown macroalgae, which primarily form aragonite. Aragonite is relatively easily dissolved in sediments, while calcite is
generally well preserved.
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747
Two key issues emerge regarding the biological flux of
CaCO3 in relation to the carbon cycle. First, it is important
to remember (Eqn 5) that the precipitation of CaCO3
increases the partial pressure of CO2 and, furthermore, the
CO2 produced per unit CaCO3 precipitated increases as
ambient CO2 increases (Frankignoule, Canon & Gattuso
1994). This is an example of a positive feedback of
increasing CO2 level on CO2 production during calcification. However, as CO2 rises, pH declines. Calcification is
highly pH sensitive, and essentially does not occur below
c. pH 7·6. Hence, while calcification can theoretically produce high concentrations of CO2, it is self regulated via
pH; this is an example of a negative feedback in a biogeochemical cycle. The factors controlling calcification in the
ocean are very poorly understood.
THE OCEANS AS CO 2 SINKS IN THE PREINDUSTRIAL PLEISTOCENE
Measurements of air trapped in ice-cores show that for at
least two thousand years prior to the Industrial Revolution,
atmospheric CO2 concentrations were approximately
280 µmol mol–1. This value appears to be typical of interglacial conditions over the last four glacial cycles. During
glacial periods, atmospheric CO2 concentrations reached
minima of approximately 180 µmol mol–1 (Jouzel et al.
1993; Raynaud et al. 1993). The causes of the
100 µmol mol–1 variations in atmospheric CO2 are unclear.
While there is a consensus that the oceans sequester CO2
during glacial periods and are a source of CO2 during
glacial–interglacial transitions, the mechanism(s) for the
exchange are contentious, especially regarding the role of
primary producers.
While terrestrial productivity declined by approximately
30% during glacial periods, there is compelling evidence
suggesting that oceanic primary productivity and the sedimentation of organic carbon were enhanced. One line of
evidence is based on the use of proxies in the sediment
which can be related to the concentration of a particular
nutrient in the overlying water column when the sediment
formed. For example, Cd in carbonates is a proxy for
HPO42– (Boyle, Slater & Edmond 1976). The 15N/14N
(δ15N) ratio in organic matter serves as a proxy for NO3–
(Altabet et al. 1995; Farrell et al. 1995; Ganeshram et al.
1995). More direct evidence related to sedimentation of
organic carbon (rather than the potential for increased primary productivity, per se) can be found in Berger,
Smetacek & Wefer (1989), Schroder (1992) Bender &
Sowers (1994) and Paytan, Kastner & Chavez (1996).
Other things being equal, enhanced sedimentation of
organic carbon during glacial periods reflects an increase
in ‘new’ primary productivity. The proxy data suggest that
the upper ocean concentrations of NO3– and HPO42– during glacial episodes were elevated. Furthermore, there is
clear evidence for increased aeolian inputs of iron (Martin
et al. 1990; Falkowski & Raven 1997). The elevated aeolian iron fluxes suggest that primary productivity in the
‘high nutrient (i.e. NO3–, HPO42–)–low chlorophyll’ parts
748
J. A. Raven and P. G. Falkowski
of the ocean would have been higher (Martin 1990;
Falkowski 1997a; Falkowski & Raven 1997; Sunda &
Huntsman 1997). The increased iron flux would have also
stimulated biological N2 fixation (Raven 1988, 1990;
Falkowski 1997a), thereby potentially enhancing the fixed
nitrogen inventory of the oceans.
The effects of these changes in NPP can be quantitatively assessed using a box modelling approach
(Broecker, Peng & Engh 1980). If the aeolian fluxes of
iron were sufficient to top up the inventory of fixed inorganic N such that the N : P ratio conformed to the
Redfield value, atmospheric CO2 would have decreased
from 280 to approximately 245 µmol mol–1 (Figs 3a & b).
Utilization of 30% of the nutrients in the Southern Ocean
would have led to a further drawdown in CO2 to
190 µmol mol–1, the glacial minimum (see legend to
Fig. 3). While these calculations do not prove that
changes in oceanic NPP were responsible for the changes
in atmospheric CO2, they clearly demonstrate their sensitivity to oceanic biological processes.
One line of evidence related to the role of primary producers in influencing glacial/interglacial CO2 concentrations in the atmosphere can be inferred from the isotopic
13 12
C/ C ratio (δ13C). If the lowering of atmospheric CO2
at the onset of a glacial episode were a result of increased
‘new’ primary productivity, atmospheric 13CO2/12CO2
would be expected to increase. This prediction is based
on the fact that marine photoautotrophs would preferentially assimilate 12C, thereby enriching the remaining
inorganic carbon with the heavier isotope. Simple diffusive equilibration of the surface ocean with the atmosphere would lead to enrichment of the latter in 13C. Two
independent lines of evidence indicate, however, that
atmospheric CO2 was enriched with 12C during the last
glacial episode. Leuenberger, Siegenthaler & Langway
(1992) showed that samples of atmospheric CO2 derived
from Antarctic ice cores had lower 13C/12C ratios during
the last glacial interval. Marino et al. (1992) used the
13 12
C/ C of the C4 terrestrial plant Atriplex confertifolia
taken from pack-rat middens of known age as a proxy for
the 13C/12C of atmospheric CO2, and also inferred a lower
13 12
C/ C in atmospheric CO2 during the last glacial
episode. These two data sets have been invoked to suggest that the oceanic ‘biological pump’ was less active in
the glacial episodes, and that the drawdown of atmospheric CO2 was a result of changes in oceanic physics
and/or physical chemistry (Keir 1992; Kerr 1992;
Leuenberger et al. 1992; Marino et al. 1992; Raven 1992;
François et al. 1997; Raven 1999).
This apparent paradox may be reconciled, albeit not yet
quantitatively, by consideration of the solubilization of
CaCO3. An enhanced flux of organic carbon to the ocean
interior would stimulate respiration. A by-product of respiration is acidification, which promotes solubilization of
CaCO3 in sedimentary particles (Archer & Maier-Reimer
1994). The liberated Ca2+ acts as a trapping agent for
HCO3–, via inorganic carbon equilibria within the ocean
and between ocean and atmosphere, account for the
Figure 3. Calculation of atmospheric CO2 between interglacials
and glacials using a simple three box model adapted from
Toggweiler & Sarmiento (1985). Figure 3(a) shows the initial
equilibrium condition during interglacials with an atmospheric CO2
partial pressure of 276 µmol mol–1, a sea surface temperature of
21·5 °C at low latitudes and 2·5 °C at high latitudes, a salinity of
34·5 kg m–3, a N : P atomic ratio of 14·7 and a phosphate
concentration of 2·15 mmol m–3 in deep water. With sufficient N2
fixation to ‘top up’ the N : P ratio to 16, other values are as shown in
Fig. 3(a) except that phosphate is lowered to 1·4 mmol m–3 in high
latitude surface water and atmospheric CO2 becomes
252 µmol mol–1. Figure 3(b) shows the situation with the N2 fixation
‘top up’ with carbonate alkalinity adjusted for the glacial temperature
(sea surface temperature 18·5 °C at low latitudes and 2·0 °C at high
latitudes) and salinity (35·9 kg m–3) values, yielding an atmospheric
CO2 partial pressure of 248 µmol mol–1. Consideration of nutrient
depletion in high latitude sea surface waters in the scenario shown in
Fig. 3(b) would yield (still with N : P = 16) a phosphate level of
2·23 mmol m–3 in deep water and 1·0 mmol m–3 in high latitude
surface waters, and an atmospheric CO2 partial pressure of
209 µmol mol–1.The 94% of the decrease in atmospheric CO2 is a
consequence of the sequestration of CO2 in the deep ocean resulting
from the enhanced export flux of carbon from the surface waters.
The model suggests that a relatively small change in N : P ratios in
the ocean can have relatively large changes in atmospheric CO2. In
this example, the change in the biological pump that accompanies
the change in N : P ratios can account for about 38% of the
difference between the glacial and interglacial atmospheric CO2
concentrations based on ice core analyses.
© 1999 Blackwell Science Ltd, Plant, Cell and Environment, 22, 741–755
Oceanic sinks for atmospheric CO2
observed decrease in atmospheric CO2 levels. This
sequence of reactions is described schematically by:
CO2 + 2H2O → (CH2O) + O2 + H2O
(primary production) (increases pH)
CH2O + O2 + H2O → CO2 + 2H2O
(respiration) (decreases pH)
CO2 + H2O + CaCO3 → 2HCO3 – + Ca2+
(calcium carbonate dissolution)
–––––––––––––––––––––––––––––––––––
(CaCO3 + H2O + CO2 → Ca2+ + 2HCO3 – )
(net reaction of above)
This carbonate intermediate pathway is quantitatively
consistent with the observed decrease in calcite preservation in marine sediments during glacial periods as well as
consistent with evidence of increased total, and exported
marine primary productivity. Such evidence includes an
increased benthic-to-planktonic gradient in δ13C, and
higher sedimentary organic carbon concentrations and
burial rates.
Further evidence which is generally consistent with this
hypothesis relating the decreased CO2 concentration in the
ocean surface-waters (and hence atmosphere) to an
increased alkalinity and pH of the ocean comes from the
δ11B work of Sanyal et al. (1995). These workers studied
the natural abundance 11B/10B ratio of B(OH)4– in fossil
foramenifera skeletons as a measure of palaeo-pH of seawater. In seawater, boron exists as boric acid (B(OH)3) and
borate ion (B(OH)4–). The δ11B ratio is 19‰ higher in
B(OH)4– than in B(OH)3 as a consequence of a thermodynamic isotope effect. The ratio of B(OH)4– : B(OH)3
increases with increasing pH. Hence a higher δ11B indicates a higher pH. The δ11B data indicated a seawater pH
that was 0·3 ± 0·1 (deep water) and 0·2 ± 0·1 (surface) units
higher during the last glaciation than it is today, which
could account for the observed difference in atmospheric
CO2 levels, albeit with some ‘implications that are difficult
to accept’ (Sanyal et al. 1995). Despite these problematic
implications, this model does predict the direction (lower,
i.e. more negative), if not necessarily the extent, of the
change in the δ13C of atmospheric CO2 in the last glacial
episode relative to today.
OCEAN INORGANIC CARBON IN RELATION
TO THE INDUSTRIAL REVOLUTION AND
GLOBAL ENVIRONMENTAL CHANGE
Over the last 250 years, the rate of fossil fuel burning and
global deforestation has increased concomitant with a rise
in human population and industrialization. These anthropogenic activities have resulted in the release of about
340 Pg of C as CO2 to the atmosphere between 1850 and
1996, of which 220 Pg resulted from the burning of fossil
fuels and 120 Pg from deforestation. The fate of this
anthropogenic CO2 is incompletely understood. The only
well-constrained sink is the atmosphere, which retains
some 42% of the total, that is, some 143 Pg C. There is
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749
considerable controversy over what happens to the remainder of the CO2 delivered to the atmosphere (Berger et al.
1989; Watson et al. 1991a,b; Falkowski & Wilson 1992;
Quay, Tilbrook & Wong 1992; Sarmiento & Sundquist
1992; Keeling 1993; Falkowski & Wilson 1993; Falkowski
1994; Frankignoule et al. 1994; Hesshaimer, Heimann &
Levin 1994; Murray et al. 1994; Sarmiento & Bender
1994; Beran 1995; Ciais et al. 1995; Bentaleb & Fortugne
1996; Bentaleb et al. 1996; Berner & Berner 1996;
Sarmiento & Quéré 1996; Doney 1997; Emerson et al.
1997; Falkowski 1997a; Fischer et al. 1997).
Clearly, in the contemporary ocean, the assumption of
steady-state does not apply (Falkowski et al. 1998).
Increasing CO2 inputs into the atmosphere from fossil fuel
burning, deforestation and cement manufacture currently
amount to some 7 Pg C per year (Table 2). The pre-industrial 0·5 Pg C per year CO2 net flux from ocean to atmosphere has been replaced by an atmosphere to ocean net
CO2 flux of up to 2 Pg C per year. The pre-industrial 0·5 Pg
C per year net CO2 flux from atmosphere to terrestrial
biota is increased to a net flux of approximately 1·5 Pg C
per year (Houghton et al. 1996; Bigg 1996). Thus, approximately 30% of the anthropogenic CO2 (i.e. some 100 Pg
C) has been taken up by the oceans between 1850 and
1996.
Most of the CO2 absorbed by the oceans is a consequence of direct solubilization due to increased partial
pressure of the gas in the atmosphere. This flux can be estimated from tracing radiocarbon released to the atmosphere
as a consequence of thermonuclear explosions and ultimately equilibrating with the ocean (Broecker & Peng
1982), as well as very high precision analyses of the concentration of total inorganic carbon in seawater over the
past several decades (Sabine, Wallace & Millero 1997).
The interaction between biological and physical processes in the ocean produces a significant
ocean–surface–atmosphere disequilibrium with respect to
CO2. The net invasion of CO2 results in the transfer of
approximately 1 Pg C per year from low-latitude to highlatitude surface ocean waters (Watson & Liss 1998; Table 2;
Fig. 2). Note that this gradient is the opposite of that prior to
the Industrial Revolution. Thus, CO2 uptake as a result of
the CO2 solubility effect is a decrease in net CO2 evolution
from the low-latitude ocean and an increased net CO2
uptake by the high latitude ocean. The rate of physicochemical uptake of CO2 by the ocean is reduced as ocean
temperature increases. Such warming has apparently
occurred over the 1850–1996 time interval (Houghton et al.
1996). The solubilization effects are complicated by any
influence that anthropogenic climate change has had on
oceanic circulation, and thus on the atmospheric CO2 transfer described (Watson & Liss 1998). While major effects on
global ocean circulation do not seem to have occurred
between the beginning of the Industrial Revolution and
today (Broecker 1997), they almost certainly will occur in
the coming centuries (Sarmiento et al. 1998).
Human activities since 1750 have increased not only
atmospheric CO2, but also the riverine and atmospheric
750
J. A. Raven and P. G. Falkowski
Flux
Pg C year-1
Volcanic/tectonic to atmosphere
From fossil fuels to atmosphere as CO2
From deforestation to atmosphere as CO2
From atmosphere in terrestrial gross primary productivity
From land biota to atmosphere by respiration and fire
From soil to atmosphere by respiration
From soil in rivers
From atmosphere to warm surface ocean
From warm surface ocean to atmosphere
From warm surface ocean to biota in gross primary productivity
From biota to warm surface ocean in respiration
From atmosphere to cold surface ocean
From cold surface ocean to atmosphere
From cold surface ocean to biota in gross primary productivity
From biota to cold surface ocean in respiration
From warm surface ocean to cold surface ocean as inorganic C
From cold surface ocean to intermediate and deep water as inorganic C
From warm surface ocean biota to intermediate and deep waters as organic
C and CaCO3
From intermediate and deep waters to warm surface ocean as inorganic C
From cold surface ocean biota to intermediate and deep waters as organic
C and CaCO3
From intermediate and deep waters to cold surface ocean as inorganic C
From biota in intermediate and deep waters to sediments as organic
C and CaCO3
0·2
5·4
2·9
102
50
50
0·7
71
71·5
21
18
20
19
10
9
19·2
50·2
2
transfer of HPO42–, NOx, NHy, organic N and transition
metals (Cornell, Rendell & Jickells 1995; Berner & Berner
1996; Falkowski & Raven 1997; Raven & Yin 1998).
Much of this additional nutrient input, with a possible
increase in global primary productivity of up to 0·1 Gt C
per annum, involves coastal waters (Walsh 1988), and so
would not have been evident in the work of Falkowski &
Wilson (1992, 1993) comparing oceanic phytoplankton
biomass and primary production over the last century. At
all events there is evidence of an increased input of potentially limiting resources to the ocean which could account
for increased organic carbon sedimentation. A less significant effect, discussed earlier, is the variable stimulation of
photosynthesis and of primary productivity in the ocean as
a result of increased CO2 levels over the last 250 years.
WHAT WILL HAPPEN WITH EXPECTED
FURTHER INCREASES IN ATMOSPHERIC CO 2
AND SURFACE OCEAN TEMPERATURE?
Given ‘business as usual scenarios’ used by the
Intergovernmental Panel on Climate Change (IPCC) to
examine climate forcing (Houghton et al. 1996) atmospheric CO2 will have doubled from the pre-Industrial
Revolution concentration by the middle of the 21st century. The solubility of CO2 in the surface ocean via the solubility pump could be partly offset by the predicted higher
temperatures. Moreover, the sequestration of dissolved
inorganic carbon in the ocean interior will probably be
reduced as a consequence of increased stratification of the
upper ocean (Sarmiento et al. 1998). Further increases in
Table 2. Carbon fluxes in the present,
industrial, part of the present interglacial.
Adapted from Fig. 1 of Watson & Liss
(1998) (our Table 1) and Siegenthaler &
Sarmiento (1993). Note that, to ensure
closure of cycles, the values differ slightly
from those in the text
22
1
31
0·2
nutrient inputs from the land to seas could increase new
production and thus, potentially, organic carbon incorporation into sediments, leading to drawdown of atmospheric
CO2 via the biological pump; however, there are caveats
regarding the efficiency and magnitude with which
increased nutrient fluxes impact on the carbon cycle.
Primary amongst these is that, despite coastal eutrophication, denitrification on continental margins often keeps
apace of nitrogen inputs, such that the net change of fixed
nitrogen in the oceans from anthropogenic sources is
extremely small, accounting for approximately 0·2 Pg C
per annum (Walsh 1991). Changes in that flux are difficult
to predict, but are not likely to lead to a large sink for
anthropogenic CO2 in the coming century. Even greater
problems ensue if atmospheric changes alter climate sufficiently to cause very markedly ‘non-linear’ responses, for
example, to ocean circulation (Broecker 1997; Falkowski
et al. 1998).
WHAT ARE THE PROSPECTS FOR
INTERVENTION IN INCREASING THE ROLE
OF THE OCEANS IN DISPOSING OF
ANTHROPOGENIC CO 2?
Background
As anthropogenic emissions of CO2 increase, there will
be increased pressure to develop methods of enhancing
sinks or reducing the source term. For the former, it has
been proposed that primary productivity could be
increased by deliberate addition of a limiting nutrient to a
part of the ocean (Martin et al. 1994). For the latter, it has
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Oceanic sinks for atmospheric CO2
been suggested that non-biological, engineering approaches
be used for trapping the CO2 from major point sources (e.g.
fossil fuel electricity generating stations) and depositing it
directly in the deep ocean, where oceanic circulation would
prevent its return to the atmosphere for decades or centuries
(Haughan & Drange 1992; Orr 1992). These two groups of
methods will be considered in turn.
Biological intervention
In principle, any of the nutrients (other than CO2!) which
limit primary productivity in different areas of the ocean
could be deliberately added to the appropriate location.
This would encompass inorganic nitrogen (NH4+ or NO3–),
HPO42–, Fe, possibly Zn and (for diatoms) Si. In practice,
the method is most conveniently applied to nutrients
required in small quantities relative to carbon (i.e. trace
elements). This approach was suggested by the late John
Martin (Martin, Gordon & Fitzwater 1990; cf. Hart 1934;
Harvey 1937) for Fe, which is a limiting resource for phytoplankton growth in two areas of the Pacific Ocean and
the Southern Ocean. There are reservations about the largescale use of this method (Fuhrman & Capone 1991). The
large increase in organic carbon export, which would follow large-scale fertilization, could lead to anoxia in the
deeper ocean. Anoxia could, in turn, enhance both
methanogenesis and denitrification. The latter would lead
to the production of N2O. Subsequent outgassing of both
methane and N2O would increase the greenhouse effect to
a greater extent than the decrease produced by the drawdown of CO2. At present atmospheric levels, a given absolute change in CH4 and N2O alters the greenhouse effect
orders of magnitude more per molecule than CO2.
With such cautionary observations in mind, small-scale
experiments have been conducted in the Eastern Tropical
Pacific, namely, IRONEX I and IRONEX II. Here FeSO4,
together with SF6 as an inert tracer, was added to 65 km2
areas of ocean to yield final concentrations of iron of
approximately 2 to 4 µmol m–3. Lessons from the first
experiment were used in designing and executing the second experiment, which yielded a very significant increase
in phytoplankton primary productivity and decrease in
ocean surface dissolved CO2 concentration (Kerr 1994;
Coale et al. 1996; Cooper, Watson & Nightingale 1996).
However, as was anticipated, the iron effect only lasted a
few weeks as a result of the oxidation and precipitation of
the iron, and its physical advection and diffusion in the
ocean. Consequently, large-scale CO2 drawdown would
need continuous Fe addition. Presumably this problem
could be overcome with engineering and technical
resources; however, estimates of increased carbon
sequestration resulting from Fe fertilization need to be
better constrained.
Non-biological intervention
The rationale of non-biological intervention is to trap the
CO2 from major point sources, such as electricity-generating
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751
stations which are major consumers of fossil fuels, and
transfer the CO2 to the deep ocean (see Haughan & Drange
1992; Orr 1992). The original suggestions involved the
deep injection of CO2 as a gas, a liquid, or a highly concentrated aqueous solution. If an appropriate site was chosen,
then the CO2 should not come back to the surface for a century or more. However, most such sites are remote from
power stations, so that the mechanics of such CO2 disposal
would be very complex (Orr 1992).
A subsequent suggestion (Haughan & Drange 1992) is
much less logistically complex. Haughan & Drange
(1992) argue that injection at approximately 150 m of
CO2 gas under pressure would convert 99% or more of
the injected CO2 into a solution of CO2 in seawater. The
solution is denser than the surrounding normal seawater
and, under appropriate conditions, could sink and thus
mimic the results of the much more complex deep injection. The prerequisites for the shallow injection procedure are that the water movement regime is adequate to
cause solution of all of the added CO2, but not so great as
to prevent sinking of the CO2-containing water.
Furthermore, even the shallow injection process has a
significant energy cost; Orr (1992) suggests that even the
removal of CO2 from the power station exhaust and its
compression would derate the station by 35% and double
its construction costs, without allowing for the costs of
movement to an appropriate shoreline.
Since injection of such quantities of CO2 would lower
the pH of seawater locally to 4·0, either deep or shallow
CO2 injection could have deleterious effects on biota near
the site of injection (Haughan & Drange 1992; Orr 1992).
In this respect, the effect would mimic the T–K boundary
bolide impact (O’Keefe & Ahrens 1989). The shallow
injection option would have more obvious impacts on benthic biota. It should be noted, however, that the CO2enriched seawater would interact with CaCO3 in sediments
with which it came into contact, thus increasing the alkalinity of the water (Orr 1992). In the extreme case, this
might even make the originally CO2-enriched deep water
act as neither a sink nor a source for atmospheric CO2
rather than a source when it is eventually upwelled, provided all of the injected CO2 has been titrated by dissolution of CaCO3 according to Eqn 7.
CONCLUSIONS
1. Oceans dominate the global C cycle over 101–106 and
more years.
2. Glacial drawdown of atmospheric CO2 in the
Pleistocene mainly involved the ocean. This was due at
least in part to increased phytoplankton primary productivity and organic C sedimentation related to (for
example) increased aeolian Fe input to the Southern
Ocean. Decreased ocean CaCO3 precipitation, with
increased soluble alkalinity, could have had a role.
3. Oceans have sequestered up to 30% of the additional
CO2 emitted to the atmosphere since the start of the
Industrial Revolution. This is due in part to the solubility
752
J. A. Raven and P. G. Falkowski
effect from increased atmospheric CO2, and in part to
increased primary productivity and organic C sedimentation.
4. Responses to further anthropogenic CO2 inputs will follow those outlined in (3), provided that global environmental change does not massively alter the oceanic
thermohaline circulation. Human intervention by Fe
fertilization of phytoplankton Fe-deficient areas of the
ocean, and direct CO2 burial in the deep ocean, could
increase the C sequestration in the ocean but with an
environmental cost.
ACKNOWLEDGMENTS
Work in J.A.R.’s laboratory on inorganic carbon acquisition by marine algae and its interaction with other environmental factors is supported by the Natural Environment
Research Council (UK) and the Scottish Office
Agriculture, Environment and Fisheries Department.
P.G.F. is supported by the National Aeronautics and Space
Administration, the US Department of Energy and the
Office of Naval Research.
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Climate-related variations in denitrification in the Arabian Sea
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