Anthropogenic Decline in High-Latitude Ocean
Carbonate by 2100
James C. Orr1, Victoria J. Fabry2, Olivier Aumont1*, Laurent Bopp1, Scott C. Doney3,
Richard M. Feely4, Anand Gnanadesikan5, Nicolas Gruber6, Akio Ishida7, Fortunat
Joos8, Robert M. Key9, Keith Lindsay10, Ernst Maier-Reimer11, Richard Matear12,
Patrick Monfray1† , Anne Mouchet13, Raymond G. Najjar14, Gian-Kasper Plattner8,
Keith B. Rodgers1*, Christopher L. Sabine4, Jorge L. Sarmiento9, Reiner Schlitzer15,
Richard D. Slater9, Ian J. Totterdell16, Marie-France Weirig15, Yasuhiro Yamanaka7, &
1LSCE, UMR CEA-CNRS, CEA Saclay, F-91191 Gif-sur-Yvette, France
2Dept. of Biol. Sciences, Cal. State University San Marcos, San Marcos, CA 92096–
3WHOI, Woods Hole, MA 02543-1543, USA.
4PMEL/NOAA Seattle, WA 98115-6349, USA.
5NOAA/GFDL, Princeton, NJ 08542, USA.
6IGPP, UCLA, Los Angeles, CA 90095-4996, USA.
7Frontier Research Center for Global Change, Yokohama 236-0001, Japan.
8Climate and Environmental Physics, University of Bern, CH-3012 Bern, Switzerland.
9AOS Program, Princeton University, Princeton, NJ 08544-0710, USA
10NCAR, Boulder, CO 80307-3000, USA.
11Max Planck Institut für Meteorologie, D-20146 Hamburg, Germany.
12CSIRO Division of Marine Research, Hobart, TAS 7001, Australia.
13Astrophysics & Geophysics Institute, University of Liege, B-4000 Liege, Belgium.
14Dept. of Meteorology, Penn. State University, University Park, PA 16802-5013, USA.
15Alfred Wegener Institute for Polar and Marine Research, D-27515 Bremerhaven,
16Southampton Oceanography Centre, Southampton SO14 3ZH, UK.
*Present address: LOCEAN, 4 Place Jussieu, F-75252 Paris cedex 05, France
†Present address: LEGOS, UMR 5566 CNES-CNRS-IRD-UPS, F-31401 Toulouse, France.
The surface ocean is everywhere saturated with respect to calcium carbonate
(CaCO3). Yet increasing atmospheric CO2 reduces ocean pH and carbonate ion
concentrations [CO32− −] and thus the level of saturation. Reduced saturation states
are expected to affect marine calcifiers even though it has been estimated that all
surface waters will remain saturated for centuries. Here we show, however, that
some surface waters will become undersaturated within decades. When
atmospheric CO2 reaches 550 ppmv, in year 2050 under the IS92a business-as-
usual scenario, Southern Ocean surface waters begin to become undersaturated
with respect to aragonite, a metastable form of CaCO3. By 2100 as atmospheric
CO2 reaches 788 ppmv, undersaturation extends throughout the entire Southern
Ocean (< 60° °S) and into the subarctic Pacific. These changes will threaten high-
latitude aragonite secreting organisms including cold-water corals, which provide
essential fish habitat, and shelled pteropods, an abundant food source for marine
Ocean uptake of CO2 will help moderate future climate change but the associated
chemistry, namely hydrolysis of CO2 in seawater, increases the hydrogen ion
concentration [H+]. Surface ocean pH is already 0.1 unit lower than preindustrial values.
By the end of the century it will probably become another 0.3 to 0.4 units lower1,2,
meaning [H+] will increase by 100 to 150%. Simultaneously, the aqueous
CO2 concentration [CO2(aq)] will increase and [CO32−] will decrease, making it harder
for marine calcifying organisms to form biogenic CaCO3. Substantial experimental
evidence indicates that calcification rates will decrease in low latitude corals3,-5, which
form reefs out of aragonite, and in phytoplankton that form their tests (shells) out of
calcite6,7, the stable form of CaCO3. Calcification rates will decline along with
[CO32−] due to its reaction with increasing concentrations of anthropogenic CO2
CO2 + CO32− + H2O →2HCO3−
even though surface waters remain supersaturated with respect to CaCO3, a condition
that previous studies have predicted will persist for hundreds of years8,4,9.
Recent predictions of future changes in surface ocean pH and carbonate chemistry
have largely focused on global average conditions1,2,10 or on the low latitudes4, where
reef building corals are abundant. Here we focus on future surface and subsurface
changes in high latitude regions where planktonic shelled pteropods are prominent
components of the upper ocean biota in the Southern Ocean, Arctic Ocean, and
subarctic Pacific11-15. Recently, it has been suggested that the cold surface waters in
such regions will begin to become undersaturated with respect to aragonite only when
atmospheric CO2 reaches 1200 ppmv, more than 4 times the preindustrial level (4×CO2)
of 280 ppmv9. In contrast, our results suggest that some polar and subpolar surface
waters will become undersaturated at ∼ 2×CO2, i.e., probably within the next 50 years.
Changes in Carbonate
We have computed modern ocean carbonate chemistry from observed alkalinity and
dissolved inorganic carbon (DIC), relying on data collected during the CO2 Survey of
the World Ocean Circulation Experiment (WOCE) and the Joint Global Ocean Flux
Study (JGOFS). These observations are centred around 1994, and have recently been
provided as a global-scale, gridded data product GLODAP16 (see Supplementary
Information). Modern surface [CO32−] varies meridionally by more than a factor of two,
from average Southern Ocean concentrations of 105 µmol kg−1 to average tropical
concentrations of 240 µmol kg−1 (Fig. 1). Low Southern Ocean [CO32−] is due to (1)
low surface temperatures and carbonate thermodynamics (high solubility) as well as (2)
large local amounts of upwelled deep water, which contain high [CO2(aq)] from organic
matter remineralisation. These two effects reinforce one another, yielding a high
positive correlation of present-day [CO32−] with temperature (e.g., R2 = 0.92 for annual
mean surface maps). Changes in [CO32−] and [CO2(aq)] are also inextricably linked to
changes in other carbonate chemistry variables (Fig. S1).
We also estimated preindustrial [CO32−] from the same data, after subtracting
data-based estimates of anthropogenic DIC17 from the modern DIC observations and
assuming that preindustrial and modern alkalinity fields were identical (see
Supplementary Information). Relative to preindustrial conditions, invasion of
anthropogenic CO2 has already reduced modern surface [CO32−] by more than 10%, i.e.,
by 29 µmol kg−1 in the tropics and 18 µmol kg−1 in the Southern Ocean. Nearly
identical results were found when instead of the data-based anthropogenic
CO2 estimates, we used simulated anthropogenic CO2, i.e., the median from 13 models
that participated in the second phase of the Ocean Carbon-Cycle Model Intercomparison
Project, OCMIP-2 (Fig. 1c).
To quantify future changes in carbonate chemistry we used simulated DIC from
ocean models that were forced by two atmospheric CO2 scenarios: the continually
increasing IPCC IS92a scenario (788 ppmv in 2100) and the IPCC S650 stabilisation
scenario (563 ppmv in 2100) (Fig. 1). Simulated perturbations in DIC relative to 1994
(the GLODAP reference year), were added to the modern DIC data; again, alkalinity
was assumed constant. To provide a measure of uncertainty we report model results as
the OCMIP median ±2σ. The median generally outperformed individual models in
OCMIP model-data comparison (Fig. S2). By year 2100 as atmospheric CO2 reaches
788 ppmv under the IS92a scenario, average tropical surface [CO32−] declines to 149
±14 µmol kg−1. This is a 45% reduction relative to the preindustrial ocean, in
agreement with previous predictions8,4. In the Southern Ocean, surface concentrations
dip to 55 ±5 µmol kg−1, i.e., 18% below the threshold where aragonite becomes
undersaturated (66 µmol kg−1).
These changes extend well below the sea surface. Throughout the Southern Ocean
(all waters south of 60°S), the entire water column becomes undersaturated with respect
to aragonite. During the 21st century under IS92a, the Southern Ocean's aragonite
saturation horizon (the limit between undersaturation and supersaturation) shoals from
its present average depth of 730 m (Fig. S3) all the way to the surface (Fig. 2).
Simultaneously, in a portion of the subarctic Pacific, the aragonite saturation horizon
shoals from depths of about 120 m to the surface. In the North Atlantic, surface waters
remain saturated with respect to aragonite, but the aragonite saturation horizon shoals
dramatically, e.g., from 2600 m to 115 m north of 50°N. The greater erosion in the
North Atlantic is due to deeper penetration and higher concentrations of anthropogenic
CO2, a tendency that is already evident in present-day data-based estimates18,17 and in
models19,20 (Figs. S4 and S5). Less spectacular changes were found for the calcite
saturation horizon. For example, in 2100 the average calcite saturation horizon in the
Southern Ocean stays below 2200 m. Nonetheless, in 2100 Weddell Sea surface waters
become slightly undersaturated with respect to calcite.
In the more conservative S650 scenario, the atmosphere reaches 2×CO2 in 2100,
50 years later than with the IS92a scenario. In 2100, average Southern Ocean surface
waters remain slightly supersaturated with respect to aragonite. But the models also
simulate that the Southern Ocean's average aragonite saturation horizon will have
shoaled from 730 m to 60 m and that the entire water column in the Weddell Sea will
have become undersaturated (Fig. 2). In the north, all surface waters remain saturated
under the S650 scenario. North of 50°N, the annual average, aragonite saturation
horizon shoals from 140 m to 70 m in the Pacific, whereas it shoals by 2000 m to 610 m
in the North Atlantic. Therefore under either scenario, the OCMIP models simulated
large changes in surface and subsurface [CO32−]. Yet these models account for only the
direct geochemical effect of increasing atmospheric CO2 because they were all forced
with prescribed modern climate conditions.
In addition to this direct geochemical effect, ocean [CO32−] is also altered by
climate variability and climate change. To quantify the added effect of future climate
change, we analysed results from three atmosphere-ocean climate models that each
included an ocean carbon cycle component (see Supplementary Information). These
three models agree that 21st century climate change will cause a general increase in
surface ocean [CO32−] (Fig. 3), mainly because most surface waters will be warmer.
Moreover, the models agree that the magnitude of this increase in [CO32−] is small,
typically counteracting less than 10% of the decrease due to the geochemical effect.
High-latitude surface waters show the smallest increases in [CO32−], even small
reductions in some cases. Therefore, our analysis suggests that physical climate change
alone will not substantially alter high-latitude surface [CO32−] during the 21st century.
Climate also varies seasonally and interannually, whereas our previous focus has
been on annual changes. To illustrate how climate variability affects surface [CO32−],
we used results from an ocean carbon cycle model forced with the daily NCEP
reanalysis fields21 over 1948-2003 (see Supplementary Information). These fields are
observationally based and vary on seasonal and interannual time scales. Simulated
interannual variability in surface ocean [CO32−] is negligible when compared with the
magnitude of the anthropogenic decline (Fig. 3b). Seasonal variability is also negligible
except in the high latitudes, where surface [CO32−] varies by about ±15 µmol kg−1 when
averaged over large regions. This is smaller than the 21st century's transient change
(e.g., ∼ 50 µmol kg−1 in the Southern Ocean). However, high-latitude surface waters do
become substantially less saturated during winter, because of cooling (higher [CO2(aq)])
and greater upwelling of DIC-enriched deep water, in agreement with previous
observations in the North Pacific22. Thus high-latitude undersaturation will be first
reached during winter.
Our predicted changes may be compared to those found in earlier studies, which
focused on surface waters in the tropics8 and in the subarctic Pacific23,22. These studies
assumed thermodynamic equilibrium between CO2 in the atmosphere and surface
waters at their in situ alkalinity, temperature, and salinity. If the pCO2 in the equilibrium
approach is taken only to represent seawater pCO2, then results agree with our non-
equilibrium approach, when the sets of carbonate chemistry constants are identical (Fig.
4). However, assuming equilibrium with the atmosphere leads to the prediction that
future undersaturation will occur too soon (at lower atmospheric CO2), mainly because
the anthropogenic transient in the ocean actually lags that in the atmosphere. For
example, with the equilibrium approach we predict that average surface waters in the
Southern Ocean become undersaturated when atmospheric pCO2 is 550 ppmv (year
2050 under IS92a), whereas our non-equilibrium approach that uses models and data
indicates that such will occur at 635 ppmv (year 2070). Despite these differences, both
approaches indicate that the Southern Ocean surface waters will probably become
undersaturated with respect to aragonite during this century. Conversely, both these
approaches disagree with a recent assessment9 that used a variant of the standard
thermodynamic equilibrium approach, where an incorrect input temperature was used
The three coupled climate-carbon models show little effect of climate change on surface
[CO32−] (Fig. 3a vs. Fig. 1) partly because air-sea CO2 exchange largely compensates
for changes in surface DIC caused by changes in marine productivity and circulation. In
subsurface waters where such compensation is lacking, these models could under- or
over-predict the extent to which [CO32−] will change as a result of changes in overlying
marine productivity (see Supplementary Information). However, we are confident in the
model-predicted trend, which only worsens the decline in subsurface [CO32−]. That is,
all coupled climate models predict a more vigorous hydrological cycle with increased
evaporation in the tropics and increased precipitation in the high latitudes24. Increased
high-latitude precipitation drives freshening of surface waters and thus greater vertical
stratification. This leads to a decrease in high-latitude nutrients, but an increase in light
availability. The latter wins out in the Southern Ocean of the IPSL-Paris model so that
at 2×CO2 there is a 10% increase in local surface carbon export of particulate organic
carbon (POC)25. Subsequent remineralisation of this exported POC within the
thermocline leads to increased DIC, which only exacerbates the decrease in high-
latitude subsurface [CO32−].
The largest uncertainty by far, and the only means to limit the future decline in
ocean [CO32−], is the atmospheric CO2 trajectory. To better characterise uncertainty due
to CO2 emissions, we compared the six illustrative IPCC SRES scenarios in the reduced
complexity, PIUB-Bern model. Under the moderate SRES B2 scenario, average
Southern Ocean surface waters in that model become undersaturated with respect to
aragonite when atmospheric CO2 reaches 600 ppmv in 2100 (Fig. 5). For the three
higher emission SRES scenarios (A1FI, A2, A1B), these waters become undersaturated
sooner (between 2058 and 2073); for the two lower emission scenarios (A1T, B1), these
waters remain slightly supersaturated in 2100. Thus if atmospheric CO2 rises above 600
ppmv, most Southern Ocean surface waters will become undersaturated with respect to
aragonite. Yet even below this level, the Southern Ocean's aragonite saturation horizon
will shoal substantially (Fig. 2). For a given atmospheric CO2 scenario, predicted
changes in surface ocean [CO32−] are much more certain than are related changes in
climate. The latter depend not only on the model response to CO2 forcing but also on
poorly constrained physical processes, such as those associated with clouds.
Ocean CO2 Uptake
With higher levels of anthropogenic CO2 and lower surface [CO32−], the change in
surface ocean DIC per unit change in atmospheric CO2 (µmol kg−1 per ppmv) will be
about 60% lower in 2100 (under IS92a) than it is today. Simultaneously, the CO32−
/CO2(aq) ratio will decrease from 4:1 to 1:1 in the Southern Ocean (Fig. 4). These
decreases are due to the well-understood anthropogenic reduction in buffer capacity26,
already accounted for in ocean carbon cycle models.
On the other hand, reduced export of CaCO3 from the high latitudes would
increase surface [CO32−], thereby increasing ocean CO2 uptake and decreasing
atmospheric CO2. Due to this effect, ocean CO2 uptake could increase by 6 to 13 Pg C
over the 21st century, based on one recent model study27 that incorporated an empirical,
CO2-dependant relationship for calcification7. Rates of calcification could decline even
further, to zero, if waters actually became undersaturated with respect to both aragonite
and calcite. We estimate that the total shutdown of high-latitude aragonite production
would lead to at most a 0.25 Pg C yr−1 increase in ocean CO2 uptake, assuming that 1
Pg C yr−1 of CaCO3 is exported globally28, that up to half of that is aragonite29,9, and
that perhaps half of all aragonite is exported from the high latitudes. The actual increase
in ocean CO2 uptake could be much lower because the aragonite fraction of the CaCO3
may be only 0.1 based on low-latitude sediment traps30, and the latitudinal distribution
of aragonite export is uncertain. Thus increased CO2 uptake from reduced export of
aragonite will provide little compensation for decreases in ocean CO2 uptake due to
reductions in buffer capacity. Of greater concern are potential biological impacts due to
The changes in seawater chemistry that we project to occur during this century could
have severe consequences for calcifying organisms, particularly shelled pteropods, the
major planktonic producers of aragonite. Pteropod population densities are high in polar
and subpolar waters. Yet only 5 species typically occur in such cold water regions and,
of these, only 1-2 species are common at the highest latitudes31. High-latitude pteropods
have 1-2 generations per year12,15,32, form integral components of food webs, and are
typically found in the upper 300 m where they may reach densities of 100's to 1000's of
individuals per m3 (refs. 11,13-15). In the Ross Sea, for example, the prominent
subpolar-polar pteropod Limacina helicina sometimes replaces krill as the dominant
zooplankton and is considered an overall indicator of ecosystem health33. In the strongly
seasonal high latitudes, sedimentation pulses of pteropods frequently occur just after
summer15,34. In the Ross Sea, pteropods account for the majority of the annual export
flux of both carbonate and organic carbon35,34. South of the Antarctic Polar Front
pteropods also dominate the export flux of CaCO336.
Pteropods may be unable to maintain shells in waters that are undersaturated with
respect to aragonite. Data from sediment traps indicate that empty pteropod shells
exhibit pitting and partial dissolution as soon as they fall below the aragonite saturation
horizon37,22,36. In vitro measurements confirm such rapid pteropod shell dissolution
rates38. New experimental evidence suggests that even shells of live pteropods dissolve
rapidly once surface waters become undersaturated with respect to aragonite9. Here we
show that when the live subarctic pteropod Clio pyramidata is subjected to
undersaturation, similar to what we predict for Southern Ocean surface waters in 2100
under IS92a, marked dissolution occurs within 48 hours at the growing edge of the shell
aperture (Fig. 6). Etch pits formed on the shell surface at the apertural margin, which is
typically ∼ 7 µm thick, as the < 1-µm exterior (prismatic layer) peeled back (Fig. 6c),
exposing the underlying aragonitic rods to dissolution. Fourteen individuals were tested.
All of them showed similar dissolution along their growing edge, even though they all
remained alive. If C. pyramidata cannot grow its protective shell, we would not expect
it to survive in waters that become undersaturated with respect to aragonite.
If the response of other high latitude pteropod species to aragonite undersaturation
is similar to that of C. pyramidata, we hypothesise that these pteropods will not be able
to adapt quickly enough to live in the undersaturated conditions that will occur over
much of the high-latitude surface ocean during the 21st century. Their distributional
ranges would then be reduced, both within the water column, disrupting vertical
migration patterns, and latitudinally, imposing a shift towards lower latitude surface
waters that remain supersaturated with respect to aragonite. At present, we do not know
if pteropod species endemic to polar regions could disappear altogether or if they can
make the transition to live in warmer, carbonate-rich waters at lower latitudes under a
different ecosystem. If pteropods are excluded from polar and subpolar regions, their
predators will be affected immediately. For instance, gymnosomes are zooplankton that
feed exclusively on shelled pteropods39,33. Pteropods also contribute to the diet of
diverse carnivorous zooplankton, mytophid and notothenoid fishes40-42, North Pacific
salmon43,44, mackerel, herring, cod, and baleen whales45.
Surface dwelling calcitic plankton such as foraminifera and coccolithophorids,
may fare better temporarily. However, the beginnings of high-latitude calcite
undersaturation will only lag that for aragonite by 50 to 100 years. The diverse benthic
calcareous organisms in high-latitude regions may also be threatened, including cold-
water corals which provide essential fish habitat46. Cold-water corals appear much less
abundant in the North Pacific than in the North Atlantic46, where the aragonite
saturation horizon is much deeper (Fig. 2). Moreover, some important taxa in Arctic and
Antarctic benthic communities secrete magnesian calcite, which can be more soluble
than aragonite. These include gorgonians46, coralline red algae, and echinoderms (sea
urchins)47. At 2×CO2, juvenile echinoderms stopped growing and produced more brittle
and fragile exoskeletons in a subtropical 6-month manipulative experiment48. For high
latitude calcifiers though, responses to reduced [CO32−] have generally not been studied.
Yet experimental evidence from lower latitude, shallow dwelling calcifiers reveals a
reduced ability to calcify with decreasing carbonate saturation state9. Given that at
2×CO2, calcification rates in some shallow dwelling calcareous organisms may decline
by up to 50%9, some calcifiers could have difficulty surviving even long enough to
experience undersaturation. Certainly, they have not experienced undersaturation for at
least the last 400,000 years49 and probably much longer50.
Changes in high-latitude seawater chemistry that will occur by the end of the
century could well alter the structure and biodiversity of polar ecosystems, impacting
multiple trophic levels. Assessing these impacts is impeded by the scarcity of relevant
Figure 1: Increasing atmospheric CO2 and decreasing surface ocean pH and
[CO32−]. a Atmospheric CO2 used to force 13 OCMIP models over the industrial
period and for two future scenarios: IS92a (I) and S650 (S). Increases in
atmospheric CO2 lead to reductions in b surface ocean pH and c surface ocean
[CO32−] (µmol kg−1). Results are given as global zonal averages for the 1994
data and the preindustrial ocean. The latter were obtained by subtracting data-
based anthropogenic DIC17 (solid) as well as by subtracting model-based
anthropogenic DIC (OCMIP median, dotted line; OCMIP range, grey shading).
Future results come from the 1994 data plus the simulated DIC perturbations to
2100 for the two scenarios; results are also shown for 2300 with S650 (thick
dashed line). The small effect of future climate change simulated by the IPSL
climate-carbon model is added as a perturbation to IS92a in 2100 (thick dotted
line); two other climate-carbon models (PIUB-Bern and CSIRO) show similar
results (Fig. 3a). The thin dashed lines indicating the [CO32−] for seawater in
equilibrium with aragonite and calcite are nearly flat, revealing weak
Figure 2: The aragonite saturation state in 2100 as indicated by ∆[CO32−]A. The
∆[CO32−]A is the in situ [CO32−] minus that for aragonite equilibrated seawater at
the same salinity, temperature, and pressure. Shown are the OCMIP-2 median
concentrations (µmol kg−1 ) in year 2100 under scenario IS92a: a surface map;
b Atlantic and c Pacific zonal averages. Thick lines indicate the aragonite
saturation horizon in 1765 (white dashes), 1994 (white solid), and 2100 (black
solid for S650 [S]; black dashes for IS92a [I]). Positive ∆[CO32−]A indicates
supersaturation; negative ∆[CO32−]A indicates undersaturation.
Figure 3: Climate-induced changes in surface [CO32−]. a The 21st century shift
in zonal mean surface ocean [CO32−] due only to climate change, from three
atmosphere-ocean climate models (CSIRO-Hobart, IPSL-Paris, and PIUB-Bern)
that each include an ocean carbon cycle component (see Supplementary
Information). b The regional-scale seasonal and interannual variability as
simulated by an ocean carbon cycle model forced with reanalysed climate
Figure 4: Key surface carbonate chemistry variables as a function of
pCO2 (µatm). Shown are both [CO32−] (solid) and [CO2(aq)] (dashed) for
average surface waters in the tropical ocean (thick), the Southern Ocean
(thickest) and the global ocean (thin). Solid and dashed lines are calculated
from the thermodynamic equilibrium approach. For comparison, open symbols
are for [CO32−] from our non-equilibrium, model-data approach vs. seawater
pCO2 (circles) and atmospheric pCO2 (squares); symbol thickness corresponds
with line thickness, which indicates the regions for area-weighted averages. The
nearly flat, thin dotted lines indicate the [CO32−] for seawater in equilibrium with
aragonite and calcite.
Figure 5: Time series of average surface [CO32−] in the Southern Ocean for the
PIUB-Bern reduced complexity model (see Fig. 3 and Supplementary
Information) under the six illustrative IPCC SRES scenarios. The results for the
SRES scenarios A1T and A2 are similar to those for the non-SRES scenarios
S650 and IS92a, respectively.
Figure 6: Shell from a live pteropod, Clio pyramidata, collected from the
subarctic Pacific and kept in water undersaturated with respect to aragonite for
48 hours. The a whole shell has superimposed white rectangles that indicate
the three corresponding magnified areas: b shell surface, which reveals etch
pits from dissolution and resulting exposure of aragonitic rods; c prismatic layer,
which has begun to peel back, increasing the surface area over which
dissolution occurs; and d aperture region, which reveals advanced shell
dissolution when compared to e a typical C. pyramidata shell not exposed to
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