The Geological Record of Ocean Acidification

Abstract

Ocean acidification may have severe consequences for marine ecosystems; however, assessing its future impact is difficult
because laboratory experiments and field observations are limited by their reduced ecologic complexity and sample period,
respectively. In contrast, the geological record contains long-term evidence for a variety of global environmental perturbations,
including ocean acidification plus their associated biotic responses. We review events exhibiting evidence for elevated atmospheric
CO2, global warming, and ocean acidification over the past ~300 million years of Earth’s history, some with contemporaneous extinction
or evolutionary turnover among marine calcifiers. Although similarities exist, no past event perfectly parallels future projections
in terms of disrupting the balance of ocean carbonate chemistry—a consequence of the unprecedented rapidity of CO2 release currently taking place.

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The Geological Record of Ocean Acidification
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The Geological Record of
Ocean Acidification
Bärbel Hönisch,
1
*Andy Ridgwell,
2
Daniela N. Schmidt,
3
Ellen Thomas,
4,5
Samantha J. Gibbs,
6
Appy Sluijs,
7
Richard Zeebe,
8
Lee Kump,
9
Rowan C. Martindale,
10
Sarah E. Greene,
2,10
Wolfgang Kiessling,
11
Justin Ries,
12
James C. Zachos,
13
Dana L. Royer,
5
Stephen Barker,
14
Thomas M. Marchitto Jr.,
15
Ryan Moyer,
16
Carles Pelejero,
17
Patrizia Ziveri,
18,19
Gavin L. Foster,
6
Branwen Williams
20
Ocean acidification may have severe consequences for marine ecosystems; however, assessing
its future impact is difficult because laboratory experiments and field observations are limited by
their reduced ecologic complexity and sample period, respectively. In contrast, the geological
record contains long-term evidence for a variety of global environmental perturbations, including
ocean acidification plus their associated biotic responses. We review events exhibiting evidence
for elevated atmospheric CO
2
, global warming, and ocean acidification over the past ~300 million
years of Earth’s history, some with contemporaneous extinction or evolutionary turnover among
marine calcifiers. Although similarities exist, no past event perfectly parallels future projections
in terms of disrupting the balance of ocean carbonate chemistry—a consequence of the
unprecedented rapidity of CO
2
release currently taking place.
The geological record is imprinted with nu-
merous examples of biotic responses to
natural perturbations in global carbon cy-
cling and climate change (Fig. 1), some of which
could have been caused by large-scale ocean
acidification. By reconstructing past changes in
marine environmental conditions, we can test hy-
potheses for the causes and effects of future-
relevant stressors such as ocean acidification on
ecosystems (1). However, for the fossil record to
be of direct utility in assessing future ecosystem
impacts, the occurrence and extent of past ocean
acidification must be unambiguously identified.
In recent years, a variety of trace-element and
isotopic tools have become available that can be
applied to infer past seawater carbonate chemis-
try. For instance, the boron isotopic composition
(d
11
B) of marine carbonates reflects changes in
seawater pH, the trace element (such as B, U, and
Zn)–to-calcium ratio of benthic and planktic for-
aminifer shells records ambient [CO
2
−
3
], and the
stable carbon isotopic composition (d
13
C) of or-
ganic molecules (alkenones) can be used to es-
timate surface ocean aqueous [CO
2
](2).
Because direct ocean geochemical proxy
observations are still relatively scarce, past ocean
acidification is often inferred from a decrease in
the accumulation and preservation of CaCO
3
in
marine sediments, potentially indicated by an in-
creased degree of fragmentation of foraminiferal
shells (3). However, it is difficult to distinguish
between the original calcification responses to
chemical changes in the surface ocean and post-
mortem conditions at the sea floor. For instance,
planktic calcifiers may secrete heavier or lighter
shells (4), but that signal may be modified at the
sea floor through dissolution or overgrowth after
deposition (5,6). This duality can introduce con-
troversy over the identification of causes and
effects, the drivers of biological change, and
REVIEW
1
Lamont-Doherty Earth Observatory of Columbia University,
Palisades, NY 10964, USA.
2
School of Geographical Sciences,
University of Bristol, Bristol BS8 1SS, UK.
3
School of Earth
Sciences,UniversityofBristol,Bristol,BS81RJ,UK.
4
Depart-
ment of Geology and Geophysics, Yale University, New Haven,
CT 06520, USA.
5
Department of Earth and Environmental
Sciences, Wesleyan University, Middletown, CT 06459, USA.
6
Ocean and Earth Science, National Oceanography Centre
Southampton, University of Southampton, Southampton SO14
3ZH, UK.
7
Department of Earth Sciences, Utrecht University,
3584 CD Utrecht, Netherlands.
8
School of Ocean and Earth
Science and Technology, Department of Oceanography, Uni-
versity of Hawaii at Manoa, Honolulu, HI 96822, USA.
9
Depart-
ment of Geosciences, Pennsylvania State University, University
Park, PA 16802, USA.
10
Department of Earth Sciences, Uni-
versity of Southern California (USC), Los Angeles, CA 90089,
USA.
11
Museum für Naturkunde at Humboldt University, 10115
Berlin, Germany.
12
Department of Marine Sciences, University
of North Carolina–Chapel Hill, NC 27599, USA.
13
Earth and
Planetary Sciences Department, University of California Santa
Cruz, CA 95064, USA.
14
School of Earth and Ocean Sciences,
Cardiff University, Cardiff CF10 3AT, UK.
15
Department of Geo-
logical Sciences and Institute of Arctic and Alpine Research,
University of Colorado, Boulder, CO 80309, USA.
16
University of
South Florida St. Petersburg, Department of Environmental
Science, Policy, and Geography, St. Petersburg, FL 33701, USA.
17
Institució Catalana de Recerca i Estudis Avançats and Depart-
ment of Marine Biology and Oceanography, Consejo Superior
de Investigaciones Científicas, 08003 Barcelona, Catalonia, Spain.
18
Institute of Environmental Science and Technology, Universitat
Autònoma de Barcelona, 01893 Barcelona, Spain.
19
Depart-
ment of Earth Sciences, Vrije Universiteit, 1081HV Amsterdam,
Netherlands.
20
W. M. Keck Science Department of Claremont
McKenna College, Pitzer College, and Scripps College, Claremont,
CA 91711, USA.
*To whom correspondence should be addressed. E-mail:
hoenisch@ldeo.columbia.edu
Organic-
walled
dinocysts
Deglaciation PETM Toarcian
OAE
Cretaceous
asteroid impact
End-Triassic
mass extinction
End-Permian
mass extinction
Calcareous
nannofossils
0 100 Time (My) 200 300
Planktic
foraminifers
Benthic
foraminifers
Shallow reef
builders
Fig. 1. Idealized diversity trajectories of the calcareous and organic fossil lineages discussed in the text.
Extinction and radiation suggest events of major environmental change throughout the past 300 My.
Calcareous plankton is shown in black, calcareous benthos in blue, and organic fossils in green, and the
line thickness indicates relative and smoothed species richness. Highlighted events (vertical red lines)
have been associated with potential ocean acidification events (Fig. 4). Calcareous organisms were not
uniformly affected at all times, suggesting the importance of synergistic environmental factors to ex-
tinction, adaptation, and evolution as well as different sensitivity due to physiological factors. Iden-
tification of a paleo-ocean acidification event therefore requires independent geochemical evidence
for ocean chemistry changes. Images of organisms are exemplary. References and further information
on the displayed organisms are available in the supporting online material.
2 MARCH 2012 VOL 335 SCIENCE www.sciencemag.org
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whether past intervals of ocean acidification
are characterized by environmental conditions
relevant for the near future. Coeval changes in
ocean circulation will also introduce regional
biases in proxy records and hence affect global
interpretations.
Here, we review the factors controlling ocean
acidification, describe evidence for the occurrence
of ocean acidification events in the past, and dis-
cuss the potential as well as weaknesses of the
geological record in helping us predict future eco-
system changes.
Is Ocean Acidification Primarily a
pH-Decline Phenomenon?
The current rate of anthropogenic CO
2
release
leads to a surface ocean environment charac-
terized not only by elevated dissolved CO
2
and
decreased pH (7) but, critically, decreased satura-
tion with respect to calcium carbonate (CaCO
3
),
a compound widely used by marine organisms
for the construction of their shells and skeletons
(8). In contrast, slower rates of CO
2
release lead
to a different balance of carbonate chemistry
changes and a smaller seawater CaCO
3
saturation
response, which may induce differential biotic
response or even no response at all, invalidating a
direct analog. The reason for a smaller saturation
response to slow CO
2
release is that the alkalinity
released by rock weathering on land must ulti-
mately be balanced by the preservation and burial
of CaCO
3
in marine sediments (Fig. 2), which
itself is controlled by the calcium carbonate sat-
uration state of the ocean (9). Hence, CaCO
3
saturation is ultimately regulated primarily by
weathering on long time scales, not atmospheric
partial pressure of CO
2
(PCO
2
). While weathering
itself is related to atmospheric PCO
2
(10), it is
related much more weakly than ocean pH, which
allows pH and CaCO
3
saturation to be almost
completely decoupled for slowly increasing at-
mospheric PCO
2
.
Using a global carbon cycle model (2), we
show the progressive coupling between CaCO
3
saturation and pH as the rate of CO
2
emissions
increases and sources (weathering) and sinks
(CaCO
3
burial) of alkalinity are no longer ba-
lanced. For rapid century-scale and thus future-
relevant increases in atmospheric PCO
2
, both
surface ocean pH and saturation state decline in
tandem (Fig. 3). The projected decrease in ocean
surface saturation state—here, with respect to
aragonite (W
aragonite
)—is an order of magnitude
larger for a rapid CO
2
increase than for a slow
[100 thousand years (ky)] CO
2
increase. Ulti-
mately, saturation recovers while the pH remains
suppressed, reflecting how changes in the oce-
anic concentrations of dissolved inorganic carbon
(DIC) and alkalinity make it possible to have
simultaneously both high CO
2
and high carbon-
ate ion concentration saturation ([CO
2
−
3
], which
controls saturation), but with the relatively greater
increase in [CO
2
] causing lower pH. The key to
unlocking the geological record of ocean acid-
ification is hence to distinguish between long-
term steady states and transient changes. We use
the term “ocean acidification event”for time in-
tervals in Earth’s history that involve both a re-
duction in ocean pH and a substantial lowering
of CaCO
3
saturation, implying a time scale on
the order of 10,000 years and shorter (Fig. 3).
Indications of Paleo-Ocean Acidification
With these criteria in mind, we review (in reverse
chronological order) the intervals in Earth’shistory
for which ocean acidification has been hypothe-
sized, along with the evidence for independent
geochemical and biotic changes. We confine this
review to the past ~300 million years (My) be-
cause the earlier Phanerozoic (and beyond) lacks
the pelagic calcifiers that not only provide key
proxy information but also create the strong deep-
sea carbonate (and hence atmospheric PCO
2
) buf-
fer that characterizes the modern Earth system
(9). Our criteria for identifying potentially future-
relevant past ocean acidification are (i) massive
CO
2
release, (ii) pH decline, and (iii) saturation
decline. We also discuss evidence for the time
scale of CO
2
release, as well as for global warming.
Events are given a similarity index that is based
on available geochemical data (table S1) and are
indicated in Fig. 4A.
Late Pleistocene deglacial transitions. The
last deglaciation is the best documented past event
associated with a substantive (30%) CO
2
rise:
189 to 265 matm between 17.8 to 11.6 ky before
pCO2(g)
CO2(aq)
+ H2OH
++ HCO3
-2H++CO
3
2-
Atmosphere [850]
Ocean [38000]
Terrestrial
biosphere [2000]
Fossil fuel CO2 emissions (8.5)
Emissions from land use change (1.0)
CO2 uptake (2.0)
Net CO2 dissolution (2.3)
Net CO2 fixation (10)
CaCO3 dissolution (0.6)
water column
CaCO3 dissolution (0.4)
sea floor
Shallow water
Corg burial (0.1)
Shallow water
CaC O3 burial (0.1)
Low temperature basaltic alteration
Carbonate (0.1)
Kerogen (0.1)
weathering
Silicate
weathering
(0.1)
Metamorphism
Volcanism
(0.1)
Calcification (1.1)
Deep sea
CaCO3 burial (0.1)
Corg oxidation (9.9)
Corg oxidation (0.1)
Reservoir inventory
values [PgC]
Processes leading to
ocean acidification
and/or reduction of
CaCO3 saturation and
their approximate
fluxes (PgC yr-1)
Processes leading to
ocean alkalinization
and/or CaCO3
saturation-increases
and their approximate
fluxes (PgC yr-1)
Volcanism
Surface sediments
Fossil fuels . . . . . .
Shales . . . . . . . . . . . . .
Mantle . . . . . . . . . . . . .
Carbonate rocks . . . .
[0.003
x
106]
[0.005
x
106]
[12
x
106]
[32
x
106]
[65
x
106]
Fig. 2. When CO
2
dissolves in seawater, it reacts with water to form carbonic
acid, which then dissociates to bicarbonate, carbonate, and hydrogen ions. The
higher concentration of hydrogen ions makes seawater acidic, but this process
is buffered on long time scales by the interplay of seawater, seafloor carbonate
sediments, and weathering on land. Shown are the major pathways of reduced
carbon (black) and of alkalinity (yellow). Processes leading to ocean acid-
ification and/or reduction of CaCO
3
saturation are indicated in red, and pro-
cesses leading to ocean alkalinization and/or CaCO
3
saturation increases are
indicated in blue. Anthropogenic perturbations are marked in italics. Ap-
proximate fluxes are printed in parentheses (PgC year
−1
), whereas reservoir
inventory values are shown in brackets [PgC]. Natural carbon cycle fluxes are
from (70); anthropogenic fluxes for 2008 are from (57), which for the land
sink is significantly above its 1990–2000 average of 2.6 PgC year
−1
due to the
2008 La Niña state (8).
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the present (B.P.) (11). Boron isotope estimates
from planktic foraminifers show a 0.15 T0.05
unit decrease in sea surface pH (12) across the
deglacial transition—an average rate of decline
of ~0.002 units per 100 years compared with the
current rate of more than 0.1 units per 100 years
(table S1). Planktic foraminiferal shell weights
decreased by 40 to 50% (4), and coccolith mass
decreased by ~25% (13). In the deep ocean,
changes in carbonate preserva-
tion (14), pH [from foraminiferal
d
11
B(15)] and [CO
2
−
3
][from
foraminiferal B/Ca and Zn/Ca
(16,17)] differed between ocean
basins, reflecting covarying
changes in deep-water circula-
tion and an internal carbon shift
within the ocean. The regional
nature of these variations high-
lights the general need for careful
evaluation of regional versus glob-
al effects in paleo-studies.
Oligocene–Pliocene. The cli-
mate of the Oligocene to Plio-
cene [34 to 2.6 million years ago
(Ma)] contains intervals of ele-
vated temperature and modest
deviations of atmospheric PCO
2
from modern values (Fig. 4). Of
particular interest has been the
Pliocene warm period [3.29 to
2.97 Ma (18,19)], which is char-
acterized by global surface tem-
peratures estimated to be ~2.5°C
higher than today (19), atmospher-
ic PCO
2
between 330 to 400 matm
(Fig. 4C) (18,20), and sea surface
pH
(T)
~0.06 to 0.11 units lower
(18) than the preindustrial. Eco-
logical responses to the warming
include migration of tropical for-
aminifer species toward the poles
(21), but there are no documented
calcification responses or increased
nannoplankton extinction rates
(22). The early to middle Miocene
(23 to 11 Ma) and Oligocene (34
to 23 Ma) were also character-
ized periods of elevated temper-
atures and slightly higher PCO
2
compared with preindustrial val-
ues (Fig. 4C) but, because of their
long duration, were not associ-
ated with changes in CaCO
3
sat-
uration (Fig. 3C).
Paleocene–Eocene. Evidence
for rapid carbon injection asso-
ciated with the Paleocene–Eocene
Thermal Maximum (PETM, 56
Ma) as well as a number of smaller
transient global warming events
(hyperthermals) during the late
Paleocene and early Eocene (58
to 51 Ma) comes primarily from
observations of large [up to –4
per mil (‰)] negative d
13
Cexcursions(23)
associated with pronounced decreases in calci-
um carbonate preservation (24). Depending on
the assumed source, rate, and magnitude of CO
2
release (25),a0.25to0.45unitdeclineinsurface
seawater pH is possible, with a reduction in mean
surface ocean aragonite saturation from W=3
downto1.5to2(1). The calcite compensation
depth (CCD) (8)roseby~2kmtoshallowerthan
1.5 km in places (24) (compared with >4 km
today). Although a pH decrease or PCO
2
increase
remains to be confirmed by geochemical proxies
for any of the hyperthermal events, the amount
of carbon injected can be modeled on the basis
of consistent carbonate d
13
C and CCD changes,
yielding between ~2000 and 6000 PgC for the
onset of the PETM (26,27). However, as with the
last glacial transition, deep sea geochemistry ap-
pears strongly modulated by regional ocean cir-
culation changes (28), which adds an additional
layer of complexity to global extrapolation and
highlights the importance of adequate spatial cov-
erage of the data.
PETM sediments record the largest extinction
among deep-sea benthic foraminifers of the past
75 My (29), and a major change in trace fossils
indicates a disruption of the macrobenthic com-
munity (30). However, the covariation of ocean
acidification, warming, and corresponding oxygen
depletion (fig. S2) (23) precludes the attribution of
this extinction to a single cause (1,29). In shallow
water environments, a gradual shift from calcar-
eous red algae and corals to larger benthic foramin-
ifers as dominant calcifiers started in the Paleocene
and was completed at the PETM with the collapse
of coralgal reefs and larger benthic foraminiferal
turnover (31). This event is recognized as one of
the four major metazoan reef crises of the past
300My(Fig.1)(32). In marginal marine settings,
coccolithophore (33) and dinoflagellate cyst (34)
assemblages display changes in species compo-
sition, but these are interpreted to reflect sensitiv-
ity to temperature, salinity stratification, and/or
nutrient availability (34,35), not necessarily acid-
ification (fig. S2). In the open ocean, the occur-
rence of deformities in some species of calcareous
nannoplankton has been described (36), but de-
spite a strong change in assemblages, there is no
bias in extinction or diversification in favor of
or against less or more calcified planktic spe-
cies (37).
Cretaceous and Cretaceous-Paleogene. The
well-known mass extinction at 65 Ma is gener-
ally accepted to have been triggered by a large
asteroid impact (38). In addition to potential ter-
restrial biomass or fossil carbon burning, the im-
pact may have caused the emission of SO
2
from
vaporized gypsum deposits at the impact site
and/or nitric acid aerosols produced by shock
heating of the atmosphere, which could have led
to acid rain and hence potentially to rapid acid-
ification of the surface ocean (38). Although
planktic calcifiers exhibited elevated rates of ex-
tinction and reduced production (22,39), reef
corals did not experience a major extinction (32),
and benthic foraminifers were not affected in ei-
ther shallow or deep waters (29). Because mul-
tiple environmental changes covaried and proxy
data for marine carbonate chemistry are not yet
available, unambiguous attribution of the planktic
extinctions to any one driver such as ocean acid-
ification is currently not possible.
The earlier Cretaceous (K) (Fig. 4A) is gen-
erally a time of massive chalk deposition (mainly
Time (years)
Atmospheric
P
CO
2
(␮atm)
300
400
500
600
1 10 100 1,000 10,000 100,000
Mean ocean
surface ⍀
aragonite
Mean ocean
surface ⍀
aragonite
2.0
Mean ocean
surface pH
SWS
7.9
8.0
8.1
8.2
8.2
Mean surface pHSWS
Time to a doubling of P
CO
2
(years)
101
102
103
104
105
3.0
4.0
8.1 8.0 7.9
2.0
3.0
4.0
A
B
C
D
CO
2
stabilization
Rising CO
2
Fig. 3. The trajectories of mean ocean surface pH and aragonite
saturation (W
aragonite
)becomeprogressivelydecoupled as the rate of
atmospheric PCO
2
change increases. The four panels show the results
of a series of experiments in an Earth system model (2). (A) Prescribed
linear increases of atmospheric PCO
2
(on a log
10
scale) from ×1 to ×2
preindustrial CO
2
, with the different model experiments spanning
a range of time scales (but experiencing the same ultimate CO
2
change). (B) Evolution of mean surface pH in response to rising CO
2
.
(C) Evolution of mean surface W
aragonite
.(D) A cross-plot illustrating
how W
aragonite
is progressively decoupled from pH as the rate of PCO
2
increase slows, with future-relevant rate of PCO
2
increase showing a
diagonal trajectory from top left to bottom right, whereas slow PCO
2
increases result in an almost horizontal trajectory toward lower pH
with very little saturation change. All plots are color-coded from red
(“fast”)toblue(“slow”). These model results include both climate
and long-term (silicate) weathering feedback. See (2)andfig.S1for
the role of these and other feedbacks.
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in the form of nannofossil calcite),
as well as one of elevated PCO
2
(Fig. 4B) and lower pH (Fig. 4D).
This association can be miscon-
ceived as evidence that marine
calcification will not be impaired
under conditions of low pH in the
future. However, this reasoning is
invalid because extended periods
of high PCO
2
(Fig. 4B) do not nec-
essarily result in a suppressed sea-
water calcite saturation state (Fig. 3)
(1,40), which exerts an impor-
tant control on organisms’calcifi-
cation (41).
Cretaceous and Jurassic oce-
anic anoxic events. The Mesozoic
oceanic anoxic events (OAEs) (in
particular, OAE 2 ~93 Ma, OAE1a
~120 Ma, and Toarcian OAE ~183
Ma) were intervals during which
the ocean’s oxygen minimum and
deep anoxic zones expanded mark-
edly (42). The onsets of these OAEs
have been linked to the emplace-
ment of large igneous provinces,
degassing large amounts of CO
2
and associated environmental con-
sequences of warming, lower oxy-
gen solubility, and possibly ocean
acidification (42). Some of the
Cretaceous OAEs were associated
with turnover in plankton commu-
nities (43). Deformities and some
minor size reduction in coccoliths,
as well as a massive increase in
the abundance of heavily calcified
nannoconids, have been observed
(44,45). However, similar to more
recent events, there is difficulty in
unequivocally attributing observa-
tions to surface water acidification
given the covariation of environ-
mental changes (46).
Because most old sea floor
(~180 Ma or older) is subducted,
the sedimentary record of the
Toarcian OAE is now restricted to former con-
tinentalmargins. Sedimentary organic and inor-
ganic carbon deposits display initially negative,
followed by positive d
13
C excursions, which is
consistent with an influx of CO
2
into the at-
mosphere followed by organic carbon burial
(42). The negative isotopic transition occurs in
distinct negative d
13
C shifts, each estimated to
occur in less than 20 ky (47) and possibly in as
little as 650 years (48). The Toarcian OAE is
associated with a reef crisis that was particularly
selective against corals and hypercalcifying
sponges (animals with a large skeletal-to–
organic biomass ratio) (Fig. 4B) (32) and with a
decrease in nannoplankton flux (49). Again,
these observations could have been a response
to any one or combination of a number of dif-
ferent contemporaneous environmental changes.
Triassic–Jurassic. The Triassic–Jurassic (T/J)
mass extinction is linked to the coeval emplace-
ment of the Central Atlantic Magmatic Province
(50). Proxy records across the T/J boundary
(~200 Ma) suggest a doubling of atmospheric
PCO
2
over as little as 20 ky (51,52), although
the absolute PCO
2
estimates differ greatly between
proxies, with leaf stomata suggesting an increase
from 700 to 2000 matm, whereas pedogenic car-
bonates indicate an increase from 2000 to 4400
matm (Fig. 4C) (2). Decreased carbonate satura-
tion is inferred from reduced pelagic carbonate
accumulation in shelf sediments (53), although
shallow water carbonate deposition can vary in
response to many parameters, not only acidifica-
tion. A calcification crisis amongst hypercalcify-
ing taxa is inferred for this period (Fig. 4B), with
reefs and scleractinian corals experiencing a near-
total collapse (32). However, the observation that
tropical species were more affected than extra-
tropical species suggests that global warming may
have been an important contributor or even dom-
inant cause of this extinction (32).
Permian–Triassic. The Permo–Triassic (P/T)
mass extinction (252.3 Ma) was the most severe
of the Phanerozoic Era and coincided, at least in
part, with one of the largest known continental
eruptions, the Siberian trap basalts. Recent es-
timates for the total CO
2
release put it at ~13,000
to 43,000 PgC in 20 to 400 ky (54–56)—an an-
nual carbon release of ~0.1 to 1 PgC [compared
with 9.9 PgC in 2008 (57)]. There is some obser-
vational evidence for carbonate dissolution in
shelf settings (54), but its interpretation is again
debated (58). There is abundant evidence for
ocean anoxia, photic zone euxinia (enrichment in
Fig. 4. Compilation of data-based
[(B) and (C)] and model-reconstructed
[(C) and (D)] indicators of global
carbon cycle evolution over the
past 300 My together with candi-
date ocean acidification events (A).
(A) Summarization of the degree to
which events (table S1) have some
similarity to modern ocean acidifica-
tion. The similarity index (table S1)
is color-coded, where red indicates
3/most similar, orange indicates
2/partly similar, and yellow indicates
1/unlike. (B)Proxy-reconstructed
atmospheric PCO
2
(2) grouped by
proxy: yellow circles indicate paleo-
sol d
13
C, light blue squares indicate
marine phytoplankton d
13
C, red
triangles indicate stomatal indices/
ratios, dark blue inverted triangles
indicate planktic foraminiferal d
11
B,
green five-pointed stars indicate
liverwort d
13
C, purple six-pointed
stars indicate sodium carbonates,
with 10-My averages shown by gray
bars. For plotting convenience, es-
timates exceeding 3000 matm are not
shown [primarily paleosol d
13
Cfrom
the uppermost Triassic/lowermost
Jurassic (2)]. (C) Ocean Mg/Ca ratios
(red triangles, left axis), reconstructed
from fluid inclusions (2) and echino-
derm fossil carbonate [red squares
(71)] together with the Phanerozoic
seawater model of (72) (red line).
Also shown (blue circles, right axis)
is [Ca
2+
] from fluid inclusions (2)
and models [blue line (72)]. (D)
Model-reconstructed changes in
mean ocean surface pH at 20-My
intervals [black line (73)].
0
2500
2000
1000
Time (millions of years before present)
1000 200 300
Carbon/climate
perturbation events
100
0200 300
10
Mg/Ca (mol mol
-1
)
[Ca
2+
] (mol mol
-1
)
K J Tr P
PgN
T
oarcian OAE
End Cretaceous
Triassic/
Jurassic
OAE 2
OAE 1a
Deglacial transitions
Cenozoic Mesozoic Paleozoic
Phanerozoic
1500
500
C
Modern (pre-industrial)
Paleocene/
Eocene
Permian/
Triassic
A
D
20
30
40
50
3000
Band estimates
up to ~5000 µatm
at 201 Ma
Atmospheric
CO
2
(␮atm)
0.0
7.4
7.6
7.8
8.0
8.2
1.0
2.0
0
4.0
5.0
6.0
7.0
Mean ocean
surface pH
SWS
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hydrogen sulfide) (59), and strong warming (54),
but no direct proxy evidence for pH or carbonate
ion changes. Knoll et al.(59) inferred the prefer-
ential survival of taxa with anatomical and phys-
iological features that should confer resilience
to reduced carbonate saturation state and hyper-
capnia (high CO
2
in blood) and preferential ex-
tinction of taxa that lacked these traits, such as
reef builders (32).
Is There a Geologic Analog for the Future?
A number of past ocean carbon-cycle perturbation
events share many of the characteristics of an-
thropogenic ocean acidification (Fig. 4 and table
S1), with the notable exception of the estimated
rates of CO
2
release. In the general absence of
direct proxy evidence for lower pH and reduced
saturation before the Pliocene, global carbon cycle
models can be used to infer the magnitude of
carbon release by fitting observed changes in the
d
13
C of calcium carbonates and organic remnants
(60). However, as well as needing information on
the source and isotopic composition of the added
carbon, the time scale of d
13
C change is critically
important to the estimation of CO
2
fluxes (25).
Because of the lack of open-ocean sediments and
increasingly poor temporal and spatial resolution
of the geological record further back in time, it is
difficult to place adequate constraints on the
duration and rate of CO
2
release. Radiometric
dating techniques are not accurate enough to
identify Mesozoic intervals of 10-ky duration,
although orbital spectral analysis of highly
resolved isotope and/or sedimentological records
can help to partly overcome this—for example,
if a d
13
C excursion is shorter or longer than one
precession cycle [21 ky (51)]. Even for the well-
studied PETM, the duration of the main phase
of this carbon injection is still debated (35,61),
and model-inferred peak rates of ≤1PgCper
year (26,61) could potentially be an underestimate.
Additional complications arise because car-
bon may not have been released at a uniform rate
and, in the extreme, may have occurred in the
form of rapid pulses. In such cases, the assump-
tion of an average emissions rate throughout
the entire duration of the pulsed release will fail
to capture the potential for episodes of intense
acidification. For instance, although the total
duration of the CO
2
release from the T/J–age
Central Atlantic Magmatic Province was esti-
mated to be ~600 ky, pulses as short as ~20 ky
have been suggested (51,62). Similarly, the main
phase of OAE1a (excluding the recovery inter-
val) was ~150 ky (45) and hence too slow for
carbonate saturation to be significantly affected
(Fig. 3), but major volcanic eruptions and thus
rapid CO
2
release could potentially have produced
future-relevant perturbations in the carbon cycle.
Substantially improved chronologies and higher-
resolution records are needed to refine estimates
of rate.
Given current knowledge of the past 300 My
of Earth’s history (Fig. 4 and table S1), the PETM
and associated hyperthermal events, the T/J, and
potentially the P/T all stand out as having excel-
lent potential as analog events, although the T/J
and P/T are much more poorly constrained be-
cause of the absence of deep-sea carbonate de-
posits. OAEs may also be relevant but were
associated with less severe volcanism (CO
2
re-
lease) than were the older events (P/T and T/J).
The last deglacial transition, although charac-
terized by temperature and CO
2
-increase, is two
orders of magnitude slower than current anthro-
pogenic change. It is also thought to largely rep-
resent a redistribution of carbon within the ocean
and to the atmosphere and terrestrial biosphere
and hence did not have as potent and globally
uniform an acidification effect as an input from
geological reserves. Because of the decoupling
between pH and saturation on long time scales
(Fig. 3), extended intervals of elevated PCO
2
such
as the middle Miocene, Oligocene, and Cretaceous
can be firmly ruled out as future-relevant analogs.
What Are the Perspectives for Using the Geological
Record to Project Global Change?
Only rapid or pulsed CO
2
release events can
provide direct future-relevant information. As-
sessment of such events critically depends on
independent geochemical quantification of the
associated changes in the carbonate system, spe-
cifically seawater-pH and CaCO
3
saturation. Geo-
chemical proxy estimates are not yet available
for the Cretaceous and beyond and need to be
obtained to verify whether ocean acidification
did indeed happen. This is challenging, because
in addition to the potential for increasing post-
depositional alteration and reduced stratigraphic
exposure, uncertainty over the chemical and iso-
topic composition of seawater increases and lim-
its our interpretation of these proxies (63,64).
Future studies will have to improve and expand
geochemical estimates and their uncertainties of
surface and deep-ocean carbonate chemistry as-
sociated with carbonate dissolution and ecolog-
ical changes. This includes finding new archives
to study the secular evolution of seawater chem-
istry but also the laboratory study of living proxy
carriers under conditions mimicking past seawater
chemistry. An unfortunate aspect of the geolog-
ical record, however, is the lack of deep-sea car-
bonates in the Early Jurassic and beyond, which
further reduces our ability to reconstruct the car-
bonate chemistry of those older events.
The sensitivity of ocean chemistry to CO
2
re-
lease, and the relationship between induced pH
and PCO
2
changes, vary through time and further
complicate the picture. For instance, seawater
calcium and magnesium ion concentrations were
different in the past (Fig. 4C). This alters the ocean’s
carbonate ion buffering capacity and hence sen-
sitivity of the Earth system to carbon perturbation
(65) because all other things being equal, higher
ambient Ca
2+
concentrations means that a lower
carbonate ion concentration is required to achieve
the same saturation and hence balance weathering.
Varying seawater Mg/Ca ratios may potentially
also affect the mineralogy of marine calcifiers,
where the more soluble high-Mg calcite predom-
inated Neogene reefs and reefs during the Per-
mian through Early Jurassic, and more resistant
low-Mg calcite predominated during the Late
Jurassic through Paleogene (66). Thus, on this
mineralogical basis the response of marine cal-
cifiers to ocean acidification and seawater geo-
chemistry during the P/T and T/J would arguably
be closer to the modern than, for example, dur-
ing the PETM (67). Improved estimates of past
seawater–Mg/Ca composition are necessary to
better evaluate all of this.
Although we have concentrated on the pros-
pects for extracting information from the geo-
logical record concerning the impact of ocean
acidification, we must question whether it really
is necessary to isolate its effect on marine orga-
nisms from other covarying factors (68). In par-
ticular, consequences of increasing atmospheric
CO
2
will also be associated with warming in the
surface ocean and a decrease in dissolved oxy-
gen concentration (69). Massive carbon release,
whether future or past, will hence share the same
combination and sign of environmental changes.
The strength of the geological record therefore
lies in revealing past coupled warming and ocean
acidification (and deoxygenation) events as an
“integrated”analog, with future and past events
sharing the same combination and sign of en-
vironmental changes. However, in additionally
driving a strong decline in calcium carbonate sat-
uration alongside pH, the current rate of (mainly
fossil fuel) CO
2
release stands out as capable of
driving a combination and magnitude of ocean
geochemical changes potentially unparalleled
in at least the last ~300 My of Earth history,
raising the possibility that we are entering an
unknown territory of marine ecosystem change.
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Acknowledgments: Funding for the “Workshop on Paleocean
Acidification and Carbon Cycle Perturbation Events”was
provided by NSF OCE 10-32374 and Past Global Changes
(PAGES). We thank the workshop participants for stimulating
discussions and contributions to this manuscript, and the USC
Wrigley Institute on Catalina Island for hosting the workshop.
Particular thanks are owed to Thorsten Kiefer of PAGES, who
initiated the workshop and supported it at all stages. This
work is a contribution to the “European Project on Ocean
Acidification”(EPOCA). Data presented in Fig. 4 are
presented in tables S2 and S3 (2).
Supporting Online Material
www.sciencemag.org/cgi/content/full/335/6072/1058/DC1
SOM Text
Figs. S1 to S3
Tables S1 to S3
References (74–217)
10.1126/science.1208277
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1
CORRECTIONS & CLarifiCations
www.sciencemag.org sCiEnCE ERRATUM POST DATE 16 MARCH 2012
Erratum
Review: “The geological record of ocean acidification” by B. Hönisch et al. (2 March,
p. 1058). The affiliation for author Carles Pelejero was incomplete. The complete affiliation
is: “Institució Catalana de Recerca i Estudis Avançats and Department of Marine Biology and
Oceanography, Institut de Ciències del Mar, Consejo Superior de Investigaciones Científicas,
08003 Barcelona, Catalonia, Spain.”
CORRECTIONS & CLarifiCations
Post date 16 March 2012
on November 19, 2012www.sciencemag.orgDownloaded from