ESDEarth System DynamicsESDEarth Syst. Dynam.2190-4987Copernicus PublicationsGöttingen, Germany10.5194/esd-9-797-2018Hazards of decreasing marine oxygen: the near-term and millennial-scale
benefits of meeting the Paris climate targetsO2 projections for multi-millennial timescalesBattagliaGiannabattaglia@climate.unibe.chhttps://orcid.org/0000-0002-6677-7969JoosFortunathttps://orcid.org/0000-0002-9483-6030Climate and Environmental Physics, Physics Institute, University of Bern, Bern, SwitzerlandOeschger Centre for Climate Change Research, University of Bern, Bern, SwitzerlandGianna Battaglia (battaglia@climate.unibe.ch)13June20189279781619October20171November201713April201815May2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://esd.copernicus.org/articles/9/797/2018/esd-9-797-2018.htmlThe full text article is available as a PDF file from https://esd.copernicus.org/articles/9/797/2018/esd-9-797-2018.pdf
Ocean deoxygenation is recognized as key ecosystem stressor of the future
ocean and associated climate-related ocean risks are relevant for current policy
decisions. In particular, benefits of reaching the ambitious
1.5 ∘C warming target mentioned by the Paris Agreement compared to
higher temperature targets are of high interest. Here, we model oceanic
oxygen, warming and their compound hazard in terms of metabolic conditions
on multi-millennial timescales for a range of equilibrium temperature
targets. Scenarios where radiative forcing is stabilized by 2300 are used
in ensemble simulations with the Bern3D Earth System Model of Intermediate
Complexity. Transiently, the global mean ocean oxygen concentration decreases
by a few percent under low forcing and by 40 % under high forcing. Deoxygenation
peaks about a thousand years after stabilization of radiative forcing and new
steady-state conditions are established after AD 8000 in our model. Hypoxic waters
expand over the next millennium and recovery is slow and remains incomplete
under high forcing. Largest transient decreases in oxygen are projected for
the deep sea. Distinct and near-linear relationships between the
equilibrium temperature response and marine O2 loss emerge. These point to
the effectiveness of the Paris climate target in reducing marine hazards and
risks. Mitigation measures are projected to reduce peak decreases in oceanic
oxygen inventory by 4.4 % ∘C-1 of avoided equilibrium warming.
In the upper ocean, the decline of a metabolic index, quantified by the ratio
of O2 supply to an organism's O2 demand, is reduced by
6.2 % ∘C-1 of avoided equilibrium warming. Definitions of peak
hypoxia demonstrate strong sensitivity to additional warming. Volumes of water
with less than 50 mmol O2 m-3, for instance, increase between
36 % and 76 % ∘C-1 of equilibrium temperature response. Our results
show that millennial-scale responses should be considered in assessments of
ocean deoxygenation and associated climate-related ocean risks. Peak hazards
occur long after stabilization of radiative forcing and new steady-state
conditions establish after AD 8000.
Introduction
Oxygen (O2) is a sparingly soluble gas and its abundance in the ocean is
decreasing under ongoing global warming . Decreasing
O2 concentrations, warming and changes in other environmental parameters
forced by anthropogenic greenhouse gas (GHG) emissions pose high risks for
marine ecosystems . The parties to the United Nations
Framework Convention on Climate Change and to the Paris
Agreement note “the importance of ensuring the integrity of all ecosystems,
including Oceans” and “the threats of irreversible damage”
Article 3. Marine changes are projected to evolve over
multi-century and millennial timescales with peak impacts occurring
potentially long after stabilization of atmospheric GHG concentrations and
peak temperatures. However, only a few studies have assessed millennial-scale impacts of
anthropogenic GHG emissions on the ocean and the reversibility of marine
changes in oxygen. Explicit quantification of the benefits of meeting the
2 or 1.5 ∘C climate targets mentioned by the Paris
Agreement with respect to the reversibility and avoidance of implied impacts
on marine oxygen and related environmental parameters, including ocean
circulation, ocean warming, metabolic viability and biological productivity
on multi-millennial timescales is missing.
Typical thresholds leading to O2 stress for many macro-organisms (hypoxia)
are around 50 mmol O2 m-3. Water with lower O2 concentrations are
effectively dead zones for many higher animals reviewed
in. Species are also sensitive to
thermal stress and their sensitivity to hypoxia increases
with higher temperatures . In the modern ocean,
oxygen-poor zones with O2< 50 mmol m-3 occupy about 5 % of its
volume . The expansion of oxygen-poor waters leads to
habitat compression, mortality and major changes in community structure where
energy preferentially flows into microbial pathways to the detriment of
higher trophic levels. Suboxic (< 5 mmol O2 m-3) or anaerobic
conditions can also lead to production of poisonous H2S within sediments
reviewed in, and decreasing O2
concentrations potentially lead to higher production and emissions of the
greenhouse gas nitrous oxide.
Observational and modeling studies
indicate an overall decline in the oceanic oxygen
content over past decades. Systematic discrepancies exist for the typically
low-oxygen tropical thermocline, where observations suggest O2 has
decreased and most models simulate increased O2 levels over the past
decades. Model projections to the end of the 21st century consistently
project the global ocean oxygen inventory to further decline with
anthropogenic climate change .
The most recent generation of Earth system models simulate global
deoxygenation by the end of the 21st century of around -1.81 % (RCP2.6) to
-3.45 % (RCP8.5; ). Impact studies have highlighted
potential habitat compression and reduced
catch potential associated with climate change at the end
of the century. Large model–model differences remain in projections of oxygen
minimum zones (OMZs; ).
Given the long residence time of anthropogenic CO2 in the atmosphere and
long equilibration timescales of the ocean overturning circulation,
anthropogenic climate change will grow and persist beyond the end of the
21st century, the typical near-term assessment timescale of climate change
. Only a few studies have assessed ocean biogeochemistry and
the oceanic oxygen content beyond this near-term timescale. Available studies
employ a range of physical and biogeochemical complexity levels from box
models to general circulation models (GCMs). Oxygen concentrations are
simulated to decline beyond the 21st century on multi-centennial timescales
. Simulations covering 2
millennia show a recovery phase thereafter .
In most studies, simulated oxygen concentrations have not
reached new steady-state conditions at the end of the simulation. Low-order
Earth system models and those of intermediate complexity (EMIC)
integrated by up to 100 000 years have demonstrated the potential for
long-term ocean oxygen depletion in response to carbon dioxide emissions and
the long equilibration timescales of ocean biogeochemical variables in
response to carbon emissions .
Multi-millennial simulations are therefore required to assess the full
amplitude of ocean biogeochemical changes and new steady-state conditions due
to anthropogenic climate change.
The distribution of O2 in the ocean results from the sum of its solubility
component set through air–sea exchange, the effect of O2 production by
phytoplankton in the euphotic zone and O2 consumption during organic
matter remineralization at depth. In modeling studies, it is possible to
identify the drivers of O2 changes by considering changes due to
solubility and changes due to oxygen consumption. When assessing the
near-term timescale at the end of 21st century, studies have shown that
different depths in the water column tend to be associated with different
dominant mechanisms of change. In the surface ocean, O2 changes tend to be
determined by changes in O2 solubility. In the subsurface, both changes in
solubility and utilization may reinforce (mid- and high latitudes) or
compensate for each other (tropics; e.g., ). In the
deep ocean, simulated O2 changes are dominated by changes in O2
utilization, which is in turn controlled by ocean ventilation see
alsofor longer timescales. Changes in the
oceanic heat content and in ocean circulation are therefore crucial for O2 changes.
Well-defined metrics that summarize the Earth system response are useful in
many aspects and may facilitate communication in the mitigation policy
context of the Paris Agreement. The Transient Climate Response to Cumulative
Carbon Emissions TCRE, or the Transient Earth System
Response to Cumulative Carbon Emissions TREX, are
such metrics. These link changes in global surface air temperature and
environmental parameters to cumulative CO2 emissions relying on near-linear relationships. Similar metrics may be developed for oceanic oxygen.
Deoxygenation is one of several marine ecosystem stressors including warming,
acidification, hypocapnia, changes in food supply and sea-level rise
. Several
key marine and coastal ecosystems may face severe risks due to climate
change even if the suggestions for low emission pathways are followed through the end of the century
. This growing body of concern
helped drive the 21st Conference of the Parties (COP21) to reach
the Paris Agreement. A goal of the Paris Agreement is to “hold the increase
in the global average temperature to well below 2 ∘C above
pre-industrial levels and pursuing efforts to limit the temperature increase
to 1.5 ∘C above pre-industrial levels” Article 2a,.
Article 4 of the Paris Agreement further states that “[in]
order to achieve the long-term temperature goal set out in Article 2, Parties
aim to reach global peaking of greenhouse gas emissions as soon as possible … and
to undertake rapid reductions thereafter in accordance with best
available science, so as to achieve a balance between anthropogenic emissions
by sources and removals by sinks of greenhouse gases in the second half of
this century” Article 4,.
In this study, we assess the effectiveness of the Paris climate targets in
reducing hazards of decreasing oceanic oxygen, ocean warming and marine
export productivity as simulated by the Bern3D Earth System Model of
Intermediate Complexity. We prescribe in the model four different, idealized
scenarios where anthropogenic GHG forcing is stabilized by AD 2300. The four
scenarios are designed to reach an equilibrium warming of 1.5, 1.9, 3.3 and
9.2 ∘C above preindustrial. Simulations are run to year AD 10 000 by
which time the ocean has reached new steady-state conditions. This allows us
to assess reversibility and the full amplitude of changes, acknowledging the
long equilibration timescale of biogeochemical variables with peak hazards
potentially occurring long after stabilization of radiative forcing in the
atmosphere. We summarize the outcomes developing global metrics which
quantify avoided marine hazards per avoided global warming on three different
time horizons. The first time horizon is the end of the 21st century, the
typical assessment timescale of climate change hazards. Here, changes in a
variable are related to changes in surface air temperature (SAT) in the year 2100. Those are contrasted to
the millennial-scale perspective where peak changes in the variable in the
course of the simulation and equilibrium changes at the end of the simulation
are related to the corresponding equilibrium warming.
In Sect. 2, we briefly describe the Bern3D model and the experimental
setup. Four different radiative forcing stabilization scenarios to meet four
temperature targets (1.5, 1.9, 3.3 and 9.2 ∘C above preindustrial)
are considered. The observation-constrained 100-member ensembles used to
explore parameter uncertainties for each scenario is introduced. In
Sect. 3, physical changes, including changes in overturning, water mass age, sea
ice, temperature, salinity and density as well as biogeochemical changes,
including changes in global oxygen inventory, the extent of oxygen minimum
zones and productivity, are presented. The compound effects of warming and
oxygen changes are assessed in the form of a metabolic index
. Underlying physical and biogeochemical processes and
mechanisms are discussed. Following earlier studies, we attribute the
contributions of O2 changes from changes in solubility, and the interplay
of ocean biology and ventilation by carrying four explicit O2 tracers and
an ideal age tracer. The graphical illustration of spatial changes is focused
on the 1.5 ∘C equilibrium warming target at the point of peak O2
decline. Additional supporting figures are given in the appendix. In
Sect. 4, the relationship between change in global mean surface air temperature (ΔSAT)
and selected impact-relevant parameters is quantified. Relationships are established for the near-term (AD 2100), the time
of the peak decline in oxygen around AD 3000 to 4000 and at year AD 10 000
when a new equilibrium has been reached in the model. Often a near-linear
relationship is found between the change in a variable of interest and the
change in SAT as simulated across the range of scenarios and ensemble members
at a distinct time. This allows us to develop new metrics to quantify avoided
marine hazards per unit change in ΔSAT at different points in time.
These quantitatively illustrate the benefits of meeting the Paris target in
terms of marine hazards. Each modeling exercise is associated with
uncertainties, and in Sect. 5, we discuss relevant uncertainties, mention
neglected processes and compare our findings to other studies. Finally, in
Sect. 6 we present implications and conclusions, and we summarize our findings
graphically for a “1.5 ∘C world” and contrast peak changes across the
range of temperature targets.
Model and simulationsBern3D
The Bern3D Earth System Model of Intermediate Complexity is a three-dimensional frictional geostrophic balance ocean model ,
which includes a sea-ice component coupled to a single-layer energy and
moisture balance model of the atmosphere and a prognostic
marine biogeochemistry module . A version with
41 × 40 horizontal grid cells and 32 vertical layers is used (see also
for model evaluation). The NCEP/NCAR
monthly wind stress climatology is prescribed at the
surface. Air–sea gas exchange, carbonate chemistry and natural Δ14C
of dissolved inorganic carbon is modeled according to OCMIP-2 protocols .
The global mean air–sea transfer rate is reduced by 19 % compared
to OCMIP-2 to match observation-based estimates of natural and bomb-produced
radiocarbon .
The biogeochemical module is based on phosphorus and simulates production and
remineralization/dissolution of organic matter, calcium carbonate and opal.
Production of particulate organic phosphorus (POP) within the euphotic zone (top
75 m) depends on temperature, light availability, phosphate and iron
following . POP remineralization within the water column
follows a power law profile . Organic matter falling on to
the sea floor is remineralized in the deepest box. Two-thirds of organic
phosphorus production form dissolved organic phosphorus (DOP), which decays with an
e-folding lifetime of 1.5 years. An updated remineralization scheme assigns
remineralization of POP and DOP to aerobic and anaerobic pathways depending
on the mean grid cell dissolved O2 concentration (see
). We introduce two power law profiles with two distinct
remineralization length scales for aerobic and anaerobic remineralization
(αaerob and αdenit). Constant
stoichiometric ratios are used for both aerobic and anaerobic
remineralization to convert biological P fluxes into carbon and alkalinity
fluxes (P : Alk : C = 1 : 17 : 117). The O2 demand for complete aerobic
remineralization is 170 molO2molPO4 and no oxygen is consumed
for anaerobic remineralization. Accordingly, aerobic remineralization in the
ocean is smaller than O2 production in the euphotic zone leading to an
O2 outgassing for steady-state conditions. The atmospheric oxygen
inventory is constant. This is justified as 99.5 % of the ocean–atmosphere
inventory is in the atmosphere and potential net fluxes of O2 from the
ocean and land to the atmosphere and fossil fuel burning have a small impact
on atmospheric O2. O2 components from O2 production, consumption and
solubility are carried as explicit model tracers to attribute changes.
Tracers add up to within 10-14 Pmol with mean inventories of 23.2, -239,
430 yielding a total of 214 Pmol, respectively (median values given). O2
components inferred from O2 saturation can result in systematic errors
from surface disequilibrium . The use of explicit tracers
avoids such systematic errors in the O2 components. As changes in the
O2 production term are small, we combine the O2 production and
consumption tracers to a O2 biology tracer when displaying sections.
We include evaluation of a metabolic index, Φ, which was proposed by
. It combines temperature and pO2 as indicators of
metabolically viable environments and is defined as the ratio of O2 supply
to an organism's resting O2 demand. We consider only relative changes
in Φ(t) relative to a reference time, t0 (average over 1870–1899):
ΔΦ(t)Φt0=pO2(t)pO2t0×expE0kB1T(t)-1Tt0-1,
where T is the absolute temperature, kB is Boltzmann's constant and the
exponential function and the parameter E0 characterize the temperature
dependence of the baseline metabolic rate. E0 only weakly affects the
relative influence of temperature and O2 gradients, and relative changes
in Φ are therefore independent of species . Here, we
consider E0= 0.87 eV (representative of Atlantic cod). For the calculation
of pO2 we pressure-correct the equilibrium constant following Eq. (5) in
. The metabolic index Φ, as proposed by
, is linear in pO2 (representing the rate of O2
supply) and decreases non-linearly with temperature (indicative of the
resting metabolic demand). One may note that the exponential curve varies
approximately linearly for typical global warming associated temperature
changes as E0/kb (≈ 10 000 K) is large.
The current set up does not include sediment interactions, temperature-dependent remineralization, variable stoichiometry, nitrogen cycle feedbacks,
atmospheric nutrient deposition, dynamic wind or freshwater input/albedo
changes from melting of continental ice sheets.
Ensemble and scenarios
To explore potential oxygen changes we set up four 100-member ensembles each
targeting a different equilibrium temperature response (∼ 1.5, 1.9,
3.3 and 9.2 ∘C above preindustrial). A feedback parameter λ
(W m-2 K-1; ), accounting for climate feedbacks that
are not explicitly treated in the Bern3D model, is chosen in combination with
radiative forcing from the Representative Concentration
Pathways RCPs; to achieve these stabilization targets. RCP2.6,
stabilizing by 2300, is run with λ values of -0.71 and
-1 W m-2 K-1,
reaching the 1.5 and 1.9 ∘C targets, respectively. RCP4.5,
stabilizing after 2100, is run with -1 W m-2 K-1 yielding a
3.3 ∘C temperature response and RCP8.5, stabilizing in the
23rd century, with -0.71 W m-2 K-1 yielding a 9.2 ∘C response
(median values given for temperature targets). Each member is spun up over
5000 years to AD 1765 boundary conditions. The radiative forcing follows RCP
scenarios RCP2.6, 4.5 and 8.5;. The RCP scenarios
are extended to year AD 10 000 by which time the ocean has equilibrated to
new steady-state conditions. Radiative forcing includes an 11-year solar
cycle up to year AD 3000. After that, all forcings are kept constant. We
employ a single-model setup, and assess uncertainties arising from organic
matter remineralization (αaerob and αdenit)
and vertical mixing (kdiff-dia). The three parameters are sampled using
the Latin hypercube sampling technique . The parameter
ranges are chosen such that all members achieve similar skill scores with
respect to observation-derived fields of natural radiocarbon
and dissolved O2 and correspond to the
values chosen in Table 1. A normal distribution is
used to sample αaerob with a standard value of -0.83 and a
standard deviation of -0.0625. αdenit is sampled uniformly
between -0.1 and -0.01. And a lognormal distribution is used to
sample kdiff-dia (standard value = 2.25 × 10-5 m2 s-1, shape parameter = 0.2,
location parameter = 0). We choose a single ensemble member with parameter
values close to the standard values as representative ensemble member to
illustrate spatial anomalies (αaerob=-0.85,
αdenit=-0.037, kdiff-dia= 2.05 × 10-5 m2 s-1).
Temporal evolution of physical variables relative to 1870–1899 for
model ensembles aiming at 1.5, 1.9, 3.3 and 9.2 ∘C warming targets.
Lines mark the median and shading marks the 90 % range of the ensemble. The
shading reflects uncertainties due to variations in the diapycnal mixing
coefficient. (a) Displays the changes in surface air and
(b) displays the changes in ocean mean temperatures.
(c,) and (d) are annual mean sea-ice areas in the respective hemisphere.
(e) Atlantic meridional overturning is the maximum of the Atlantic
meridional overturning stream function below 400 m depth and
(f) Indo-Pacific meridional overturning is the minimum of the
Indo-Pacific meridional overturning stream function below 400 m depth.
Pre-industrial characteristics
The ensemble produces a range in overturning strengths, remineralization
fluxes and O2 distributions. The following numbers represent the 90 %
confidence ranges of important model characteristics across the ensemble. The
maximum of the Atlantic meridional overturning circulation (AMOC) stream function below 400 m
depth ranges from 16.5 to 19.7 Sv. The minimum of the Indo-Pacific
meridional overturning stream function below 400 m depth (Indo-Pacific MOC)
ranges between -13.6 and -15.6 Sv. Export of particulate organic matter at
75 m ranges from 9.0 to 11.4 Gt C yr-1. The simulated oxygen inventory
ranges between 195 and 230 Pmol given the three parameters and the simulated
oxygen distribution covers the observational range and spatial pattern well
see Figs. 3 and 7a and Table D1 of. Biases exist in
the simulated extent of OMZs. The volume of suboxic conditions (O2< 5 mmol m-3)
is overestimated by a factor of 5 but water column
denitrification fluxes are well within current estimates Table D.1.,
Fig. 2c in. This is a common model bias in EMICs and GCMs
. Vastly enhanced spatial resolution
may be required to simulate equatorial physics and ecosystems in better
agreement with observations .
Marine changes in temperature, circulation and biogeochemistry
We first describe the evolution of important physical quantities that impact
O2 concentrations. Figure displays the temporal changes
in global mean surface air and ocean temperature, the evolution of annual
mean sea-ice area in the Northern and Southern hemispheres and the Atlantic
and Indo-Pacific meridional overturning circulations.
Temporal evolution of critical variables relative to 1870–1899 for
model ensembles aiming at 1.5, 1.9, 3.3 and 9.2 ∘C warming targets.
Lines mark the median and shading marks the 90 % range of the ensemble. The
shading reflects uncertainties due to variations in the diapycnal mixing
coefficient and the aerobic and anaerobic remineralization length scales of
particulate organic matter (αaerob and
αdenit). (a) is the change in the total oceanic
O2 inventory. (b)O2 solubility is the explicitly
traced solubility component of oceanic oxygen, (c) is the explicit
O2 production tracer, and (e) is the explicit O2
consumption tracer. (h) Oxygen-poor waters are taken as the volume of
water with O2< 50 mmol m-3.
Changes in O2 and its components at time of peak O2 decline
(AD 3150) relative to preindustrial steady state for a single representative
ensemble member reaching the 1.5 ∘C warming target. (a) Change in total
O2, (b) change in O2 due to biology, (c) change
in O2 due to solubility, (d) change in ideal age.
In response to the RCP scenarios, atmospheric temperatures rise and stabilize
after ∼ 1000 years (Fig. a). The four ensembles
reaching 1.5, 1.9, 3.3 and 9.2 ∘C above preindustrial surface air temperature
show an equilibrium ocean warming of 1.1, 1.3, 2.0 and 5.5 ∘C,
respectively (median values given). Sea ice retreats in both hemispheres
(Fig. c and d). The retreat is more pronounced for higher
forcing. In the Southern Hemisphere, even the lower forcing levels show
strong decline in the annual mean sea-ice area and sea ice vanishes for
higher forcing. The warming perturbation causes the AMOC and Indo-Pacific MOC
to decline transiently (Figs. e, f and ).
The larger the forcing and implied changes in
stratification, the larger the peak decline in overturning (Fig. e
and f). The decline is likely driven by upper ocean warming,
leading to increasing surface-to-deep density gradients as further modulated
by salinity changes (Fig. ). The deep ocean water mass age
increases in response to the slowed overturning (Figs. d
and d). As retreating sea-ice increases wind stress over these
newly exposed areas, younger water masses form in the upper ocean of the
Southern Ocean (Fig. d). Reduced convection may contribute
to a younger upper ocean through decreased entrainment of old deep water (not
quantified within the scope of this paper). As the model tends to equilibrate
under the sustained radiative forcing, the surface-to-deep gradients in the
density anomalies diminish (Fig. ), the meridional overturning
circulation recovers (Fig. e and f) and anomalies in water mass
age become again smaller (Fig. d versus Fig. ).
The final circulation state is close but not
identical to the preindustrial steady-state circulation. Maximum overturning
strength in AMOC and the Indo-Pacific MOC varies by less than ±1 Sv
around the initial value. At the new steady state, the maximum in AMOC below
400 m tends to be lower under higher forcing, whereas the maximum in the
Indo-Pacific MOC below 400 m tends to be higher under higher forcing
(Fig. e and f). It is difficult and beyond the scope of this paper to
conclusively explain such subtle changes in ocean dynamics and overturning
(Fig. ), likely linked to the complex changes in density
(Fig. ) and sea-ice retreat (Fig. c and d). However,
these differences have direct consequences for the projected global water
mass age and by that for oceanic oxygen (Fig. ) at the new
equilibrium as further discussed below.
The response in oceanic oxygen is complex and characterized by an initial
decline followed by a recovery phase (Fig. a). In line with
earlier studies ,
our results demonstrate the potential for large changes in
marine oxygen under anthropogenic forcing, a large inertia in the response
and a slow and partially incomplete recovery of the perturbation.
Transiently, the whole ocean oxygen inventory decreases by a few percent
(6 %) under low forcing and by as much as 40 % under high forcing (median
values given). The minimum in oxygen occurs about a thousand years after
stabilization of radiative forcing, and it takes several millennia to
approach a new equilibrium. Then, the global ocean O2 inventory is a few
percent higher than at preindustrial conditions under low and intermediate
forcing and remains depleted by around 8 % in the high forcing case.
Figure further explains the temporal evolution and interplay
of the underlying drivers. In all cases, the changes in global oxygen
inventory (Fig. a) strongly correlate with water mass age
(Fig. d) and are also impacted by gradual oxygen loss due to
warming as evidenced by the evolution of the O2 solubility tracer
(Fig. b). Inventory changes based on the O2 production
tracer (Fig. c) are negligible; changes equilibrate with the
atmosphere and only a small fraction remains in the ocean. The O2
consumption tracer (Fig. e) determines the shape of the global
O2 signal (Fig. a). Its decline and recovery phase is
strongly correlated with the evolution of ideal age (Fig. e
and d, see also Fig. b and d). As has been shown previously
e.g.,, the high correlation between changes
in O2 and ideal age and the absence of a direct relationship between
changes in remineralization fluxes and O2, indicate that circulation
changes are the major contributors to changes in O2. Changes from
remineralization fluxes include both changes in absolute aerobic
remineralization fluxes and changes in the relative share of denitrification
(Fig. i). An increased share of denitrification at organic
matter remineralization, for instance, effectively constitutes an implicit
O2 gain. Denitrification fluxes correlate with the volumetric expansion of
OMZs and are also impacted by changes in remineralization fluxes within them
(Fig. i). The recovery level of the O2 consumption tracer
(Fig. e) reflects the global recovery level of ideal age
(Fig. d), where younger water masses are associated with less O2
consumption and therefore higher O2 concentrations. The total O2
recovery level (Fig. a), on the other hand, is diminished due
to O2 loss from solubility (Fig. b). As such, 1.5 to
3.3 ∘C warming targets reach similar global O2 equilibrium levels
for different reasons. The 1.9 and 3.3 ∘C warming targets tend to
result in younger water masses, which would increase O2 due to less O2
consumption compared to 1.5 ∘C warming targets. As those scenarios are
also associated with higher warming, they lose more O2 due to less
solubility compared to 1.5 ∘C warming targets and yield similar global
anomalies despite more pronounced spatial patterns. The 9.2 ∘C warming
target reaches a lower equilibrium O2 inventory compared to preindustrial
due to high O2 loss from solubility (-44.1 Pmol).
Changes in potential ecosystem stressors at peak O2 decline
(AD 3150) relative to preindustrial steady state for a single representative
ensemble member reaching the 1.5 ∘C warming target. Results are
displayed for a cross section through the Atlantic (25∘ W), across the
Southern Ocean (58∘ S) and into the Pacific (175∘ W). Changes
in POM export at 75 m (c) and in surface PO4 concentrations (d)
are displayed along the same section.
We illustrate spatial changes in critical variables for a single
representative ensemble member (see Sect. 2.2) at its peak O2
decline,
which occurs at year AD 3150 and amounts to 5 % (Fig. ).
The member eventually reaches the 1.5 ∘C warming target.
Figure displays anomalies in total O2
(Fig. a), and the contributions from biologically mediated
changes (termed “biology” below, Fig. b) combining the
changes in the O2 production and consumption tracer and from changes in
solubility (Fig. c). In the upper ocean O2
concentrations tend to increase due to biology and decrease due to
solubility. Such compensating mechanisms have been documented elsewhere
e.g.,. The resulting changes in O2 are less
pronounced than the changes in each component. The increase in O2 due to
biology stems from younger water masses and less export in the low and mid-latitudes (see next paragraph and Fig. c). O2 changes
show strong spatial correlation with changes in water mass age
(Fig. a and d). Largest decreases in O2 are simulated in
bottom waters in line with older water mass age. The equilibrium response in
O2 for this 1.5 ∘C warming case is characterized by slight O2
decreases in the Atlantic, caused mainly by less solubility, and increases in
the Southern Ocean and deep Pacific, caused by higher overturning and less
sea-ice coverage in the Southern Ocean compared to preindustrial (Fig. ).
Global export production is simulated to decline over the first few
centuries, and reach higher values under new steady-state conditions
(Fig. g). The decline is stronger for higher forcing, while
the recovery level of global export production is similar across the
scenarios. Bern3D transiently simulates decreased export in the mid- and low
latitudes (Fig. c; see also )
as a result of increased stratification (Fig. c, f and i)
and reduced nutrient concentrations in the surface ocean
(Fig. b). In the high latitudes, the model simulates
increased export production, as a result of less temperature and light
limitation as surface waters warm and sea ice retreats. This pattern of
decreased export in mid- and low latitudes and increased export in high
latitudes is similar across the scenarios. Export production in the low
latitudes fully recovers for lower forcing and partially recovers for higher
forcing. The lower recovery level in the low latitudes is compensated by
higher increases in the high latitudes for high forcing. The magnitude of
positive and negative changes increases with forcing, but the global
anomalies remain comparable at the end of the simulation.
Next to changes in export, we consider the evolution of a metabolic index in
the upper ocean which integrates effects of changes in O2 and temperature
at the organism level (Figs. f and e).
The globally averaged, upper ocean (depth < 400 m) metabolic index declines
throughout the simulation dominated by increased temperatures
(Fig. f). The metabolic index, Φ, decreases
in most places in line with warming and lower pO2 (Fig. a,
d and e). The O2 gain in upper ocean waters is able to
counteract the adverse effect of warming in some high-latitude environments.
In other places with higher pO2, the temperature increase dominates the
response in Φ. Near-bottom waters in the Pacific are prone to largest
reductions in Φ, driven by large decreases in pO2 (Fig. e).
Oxygen-poor waters (O2< 50 mmol m-3, Fig. h) are
simulated to transiently increase across all scenarios. The response is
characterized by high uncertainty as introduced by the sampled parameters.
Under new equilibrium conditions, the volume of low O2 waters is reduced
for low and intermediate forcing and remains higher than pre-industrial levels in
the high forcing case.
Changes in marine ecosystem stressors versus changes in global mean
surface air temperature at three distinct points in time relative to 1870–1899.
The colored dots indicate results of the four 100-member ensembles
targeted to reach 1.5 (blue dots), 1.9 (light blue dots), 3.3 (orange dots)
and 9.2 ∘C (red dots) warming targets. The lines connect results of
individual ensemble members at AD 2100 (gray), at time of peak decline of
each variable (light blue) and by the end of the simulation (light green)
when a new equilibrium state has been reached. Peak and equilibrium changes
in variables of interest are related to the corresponding equilibrium
temperature response, while changes at the end of the 21st century are
related to the transiently realized warming. Peak and equilibrium responses
are indistinguishable in the figure for the metabolic index (d) and
the ocean temperature change (e). Speak is the peak
sensitivity of each variable per ∘C equilibrium warming.
Turning to uncertainties in our perturbed parameter ensemble, we find that
variations in the vertical diffusion parameter (kdiff-dia) dominate the
uncertainty in the globally averaged evolution of ideal age, sea-ice cover,
temperature and O2. The modeled uncertainty in the volume of low O2
waters is dominated by different values of the αaerob parameter.
Whether a threshold in O2 concentration is met depends on the
pre-industrial tracer distribution. Longer remineralization length scales
bring more remineralization to depth, leading to higher O2 consumption.
Metrics linking global warming to marine hazards
The purpose of this section is to quantify the relationship between changes
in global mean surface air temperature (SAT), the target variable of the
Paris Agreement, with selected aggregated metrics for marine ecosystem
stressors. In this way, we link marine hazards to the temperature target of
the Paris Agreement and quantify avoided marine change per unit of avoided
global warming. Specifically, we investigate the relationship of SAT with
changes in the marine O2 inventory, ocean temperature and the metabolic
index of , the volume occupied by hypoxic water and in low-latitude export production (30∘ S–30∘ N) across the range
of warming scenarios in our ensemble (Fig. ). Distinct
and often near-linear relationships emerge. Near-linearity allows us to
characterize the benefits of avoided warming by single sensitivities,
corresponding to the slopes of the relationships displayed in Fig. .
The relationships between SAT and marine hazard metrics critically depend on
the time horizon considered (Fig. ). We compare and
contrast changes at the end of the 21st century, the typical assessment
timescale of climate change, to peak and equilibrium changes at the
millennial timescale. Peak and equilibrium changes are related to the
corresponding equilibrium temperature response, while changes at the end of
the 21st century are related to the transiently realized warming at the end
of the 21st century. Larger magnitudes are simulated on millennial timescales
compared to the near-term end of the 21st century. Assessment of ocean
deoxygenation by the end of the 21st century, therefore, underestimates the
full amplitude of change.
Transient (end of 21st century), peak (AD ∼ 3000) and equilibrium
(AD ∼ 8000) oxygen changes exhibit distinct relationships with their
corresponding warming (Fig. a). At the end of the
21st century, simulated oxygen decreases by 0.68 % ∘C-1 of realized
warming (median values). At peak oxygen decline, this sensitivity increases
and oxygen decreases by 4.4 % ∘C-1 of equilibrium temperature
response. In other words, an avoided warming of 1 ∘C, avoids a peak
decline in marine O2 inventory of 4.4 %. The linear relationship breaks
down for the equilibrium response. While 1.5 to 3.3 ∘C warming targets
lead to similar and higher oxygen levels, the 9.2 ∘C warming target
results in lower oxygen levels compared to preindustrial as discussed in the
previous section. The relationships generally hold across the sampled parameter space.
The volume of low-oxygen waters is particularly sensitive to warming and
parameter uncertainty (Fig. b). We illustrate the
sensitivities at the example of the volume of waters with O2< 50 mmol m-3.
At the end of the 21st century, there is a 1.7 % increase in
this volume per ∘C of realized warming. Peak increases scale with
63 % ∘C-1 of equilibrium temperature response. Uncertainties in
remineralization cause a spread in this response ranging from
36 to 76 % ∘C-1 of equilibrium temperature response (90 %
confidence range): the longer the remineralization length scale, the higher
this sensitivity. Pre-existing low O2 waters expand and new low O2
waters may develop in near-bottom environments for higher forcing levels.
While the lower temperature targets yield lower volumes of low-oxygen waters,
the 9.2 ∘C target yields higher low O2 volumes under new steady-state conditions. In brief, hypoxic waters expand over the next millennium
across the scenario range and recovery towards modern conditions is slow and,
in the case of high forcing, incomplete. Acknowledging millennial timescales,
the hazard of expanding low O2 waters is much larger than when assessed on
the near-term timescale.
The sensitivity of changes in low-latitude export production (30∘ S–30∘ N)
is similar at the end of the 21st century and at time of peak
decline. Changes scale with 2.2 % ∘C-1. Export production in the
low latitudes recovers for lower forcing, but remains reduced in the high forcing case.
Decreases in the metabolic index of the upper ocean scale linearly with
forcing: 5.1 % ∘C-1 of realized warming at the end of the
21st century and 6.2 % ∘C-1 of equilibrium temperature response.
Likewise, global mean oceanic temperatures increase by
0.099 ∘C per degree of realized warming at the end of the
21st century and by 0.56 ∘C per degree of warming at equilibrium. In conclusion,
the compound hazards related to deoxygenation and warming, as indicated by
the metabolic stress index, evolve over millennia and increase with
increasing anthropogenic forcing and with time.
The magnitude and intensity, as well as the duration of oxygen-related
transient hazards, and thus the severity of the hazards, increase with
increasing temperature targets. The severity combines magnitude and duration
of a hazard into one quantity. It may be defined as the time integral of a
hazard. The severity of the hazard of expanding hypoxic waters, for example,
corresponds to the area under the scenario curve shown in
Fig. h (the area enclosed by the null line and the modeled
evolution, here until the end of the simulation). Figure a, h,
and g illustrate that the severity of the three hazards of decreasing mean
oxygen concentration, expanding hypoxic waters and reduced export of
particulate organic matter providing food for deep sea organisms increases
strongly from low to high temperature targets.
Uncertainties in O2 projections
The pattern and magnitude of simulated global O2 changes are determined by
the response of the overturning circulation. O2 loss due to less O2
solubility at higher temperatures also gradually decreases oceanic O2. Only a few multi-millennial simulations with GCMs currently exist.
The response of the overturning circulation on long timescales differs among
available model simulations (including EMICs and GCMs). Uncertainties in the
equilibrium climate sensitivity additionally impact projections of O2 loss
due to solubility. These uncertainties directly impact projections of oceanic oxygen.
Similar circulation dynamics as simulated here (Fig. e and f)
were found by based on EMIC simulations over 10 000 years
with ECBILT-CLIO, which features a dynamic, quasi-geostrophic
atmosphere. also found similar AMOC and Indo-Pacific
MOC characteristics for their EMIC UVic 2.7, which includes an atmospheric
energy balance model with fixed wind fields similar to the Bern3D model, over
a 2000-year simulation. , on the other hand, found
different overturning characteristics in a simulation with a state-of-the art
Earth system model (MIROC 3.2 for a 4 ×CO2) over 2000 years. There AMOC
slowed down with no recovery, while Antarctic Bottom Water formation decreased only slightly and
gradually increased thereafter. Predictions of AMOC have received more
attention so far, and AMOC slowdown and partial or full recovery emerges in
other multi-millennial simulations .
AMOC and Southern Ocean overturning in CMIP5 Earth system models were analyzed
by . They found AMOC and Southern Ocean overturning is
positively correlated in most CMIP5 models by the end of the 21st century.
Generally, preindustrial circulation states, magnitudes and timing of changes
are highly model and scenario dependent such that the long-term evolution of
meridional overturning is uncertain. As oxygen changes are dominated by
circulation changes, this makes the oxygen prediction highly model and
scenario dependent as well. The simulated timing and strength of peak O2
decrease in Bern3D is similar to what
found (AD 3000, 30 % for SRES A2
high emission scenario/SAT ∼ 10 ∘C in UVic 2.7). Other comparable simulations show earlier peaks and smaller magnitudes
such as AD 2600, 16 % decrease for RCP8.5/ΔSAT ∼ 7 ∘C in
CLIMBER-3α in , and after 800 model years, 10 % for
4 ×CO2/ΔSAT ∼ 8.5 ∘C in MIROC 3.2 in .
Major physical limitations of our simulations concern prescribed winds and
ice sheets. Future model studies may include sensitivity simulations with
prescribed changes in the wind stress over the ocean
e.g., and prescribed meltwater fluxes or apply Earth
system models with interactive atmospheric dynamics and ice sheets. Our
study, as is the case for most climate change simulations, does not include
melting of continental ice sheets, which would tend to further (transiently)
reduce circulation and increase the equilibrium climate sensitivity.
Current generation GCMs, such as is the case for Bern3D, have difficulty
simulating the current distribution of OMZs due to missing physical processes
operating at small spatial scales, such as eddies and zonal jets
or missing biogeochemical characteristics. Large
model–data and model–model discrepancies exist .
recently achieved improved representation of OMZs
introducing temperature and oxygen dependence of the remineralization profile
within a GCM (GFDL ESM2M). In our ensemble, the magnitude of peak increases
in low O2 waters depend strongly on the rate of organic matter
remineralization. Temperature-dependent feedback mechanisms, neglected here,
may be addressed in future studies. Both particulate sinking speed and local
remineralization rates, which control the remineralization profile, have been
shown to be sensitive to temperature. While higher temperatures increase
bacterial activity and therefore remineralization they
decrease viscosity and therefore increase sinking speed .
The net effect on the remineralization profile is correspondingly uncertain.
In addition, ecosystem structure influences the size and density of organic
particles available for export . Given
these existing uncertainties and the coarse-resolution physical models, the
projections of OMZs has to be viewed with caution. Despite simulated lower
background concentrations of O2 in the subsurface ocean, the volumes of
low O2 waters decrease for steady-state conditions in the model.
We neglect a number of biogeochemical feedback mechanisms that could alter
biological productivity in the surface ocean and thus
remineralization fluxes in the water column. Any mechanisms that would
increase remineralization would tend to decrease oceanic oxygen, and
mechanisms that decrease remineralization would increase the oceanic oxygen
content. Future studies may address feedbacks from sediment interactions and
imbalances from riverine input and burial such as, temperature-dependent remineralization and variable
stoichiometry. Further investigations may also address nitrogen cycle
dynamics and assess the interplay of denitrification and N fixation and of
external atmospheric and terrestrial nitrogen sources.
Implications and conclusion
In Bern3D, strong deoxygenation in all basins is projected to peak long after
the end of the 21st century, and new steady-state conditions establish after
AD 8000 in scenarios where radiative forcing is stabilized in the next
century. The equilibration timescale of oceanic oxygen is therefore longer
than the thermal equilibration timescale of both the atmosphere (∼ 1000 years)
and the ocean (∼ 4000 years). Based on CMIP5 models,
discuss the deep sea ecosystem implications of climate
change by 2100. Deep sea ecosystems provide a range of services including habitat
provision, nursery grounds, trophic support and refugia to biodiversity
reviewed in. Biogeochemical changes such as
deoxygenation, warming, acidification and less food availability will likely
be accompanied by exploitation of mineral resources, overfishing and dumping
of pollutants and microplastics. We project the largest biogeochemical changes
to occur after 2100 and to aggravate over millennia. How these changes will affect
deep sea ecosystems is poorly understood, but the adaptation to stress may be
limited by their slow growth rates and long generation times .
Contrasting hazards of ecosystem impacts expressed in definitions of
hypoxia, metabolic viability of the upper ocean and food availability.
(a) Changes for a 1.5 ∘C warmer world at present, at the end
of the century and compared to peak changes. (b) Peak changes across 1.5,
1.9 and 3.3 ∘C temperature targets. Lines correspond to the median
response across each ensemble relative to 1870–1899.
Figure a contrasts near-term (AD 2100) and peak changes
(relative to 1870–1899) in definitions of metabolically viable habitats in the
upper ocean, hypoxia and food availability as projected by Bern3D for a
1.5 ∘C warmer world. Export in low latitudes (30∘ S–30∘ N)
as an indicator of food availability is reduced by maximally
4 % over the course of the simulation in this scenario. Median decreases in
the metabolic index, representing viable habitat reductions of the upper
ocean, amount to 11 % for a 1.5∘ warmer world. The volume of low-oxygen waters is particularly sensitive to anthropogenic warming and peak
changes occur after the end of the 21st century. The volume of water with
O2< 50 mmol m-3 changes by 6.6 % by the end of the century and by
14 % at its peak. Meeting the 1.5 ∘C climate target of the Paris
Agreement requires very fast and very stringent emission reductions
. Estimates by
for a range of scenarios show that post-2017 allowable
carbon emissions from fossil fuel need to be lower than 320 Gt C to meet the
1.5 ∘C target with 66 % probability (320 Gt C is derived by
adjusting the emission limit as displayed in Fig. 4 of
, for year 2000 using fossil emissions published by the Global Carbon Project).
This corresponds in the most optimistic scenario to only slightly more than
3 decades of current fossil fuel use. The nationally determined
contributions, outlining emission mitigation actions by the parties of the
Paris Agreement, need to be more ambitious and broader in scope to meet the
1.5 or 2 ∘C targets . Such efforts
would lead not only to lower warming compared to the current emission
trajectory, but also have the benefit of reduced marine hazards as
investigated here (Fig. b).
Higher temperature targets increase the hazard of ecosystem impacts as
expressed in the investigated variables. In particular, definitions of peak
hypoxia exhibit strong sensitivity to additional warming (Fig. b).
Definitions of deoxygenation, marine food scarcity and marine
aerobic habitat reduction are aggravated for the 2 ∘C compared to the
1.5 ∘C temperature target and investigated hazards are strongly
amplified in a world where surface air temperature is stabilized at
3.3 ∘C (Fig. b). Unbounded use of carbon emissions from
existing fossil resources is projected not only to lead to a global warming
on the order of 10 ∘C Figs. a and ;,
but also to a peak reduction in global mean oxygen inventory by almost a factor
of 2 (Figs. a and a).
We find a near-linear relationship between impact-relevant marine hazards
and global mean surface air temperature. This allows us to quantify avoided
hazards per unit of avoided global warming. For example, emission mitigation
measures would help to reduce peak O2 loss by 4.4 % ∘C-1 of
avoided equilibrium warming.
The Earth system response timescale to climate change spans several millennia
such that anthropogenic perturbations to greenhouse gas concentrations commit
the Earth system to long-term, irreversible climate change .
Our simulations show that the long-term fate of oceanic oxygen is
characterized by an initial decline followed by a recovery phase. Peak
decline and associated potential adverse ecosystem impacts are projected long
after stabilization of radiative forcing in the atmosphere. This adds to the
list of long-term Earth system commitments including warming, acidification
and sea-level rise assessed elsewhere .
Long-term, multi-millennial perspectives are thus required for a full account
of climate-related ocean risks.
Model output is available upon request to the corresponding
author (battaglia@climate.unibe.ch).
Spatial properties of a representative ensemble member reaching the 1.5 ∘C warming target
In this appendix we document additional spatial properties of the
representative ensemble member reaching the 1.5 ∘C warming target.
Figure illustrates the meridional overturning
stream function for pre-industrial (PI), year AD 2050, where the AMOC is at its lowest value,
and for new steady-state conditions. Figure illustrates the
evolution of temperature, salinity and density anomalies across a transect
from the Atlantic Ocean, through the Southern Ocean and into the Pacific at
AD 2100 and AD 3150, when the O2 inventory values are at their lowest, and
for new steady-state conditions. Figure shows the
anomalies in total O2 and the contributions from biology and solubility
components for new steady conditions relative to PI. In addition, anomalies
in ideal age are shown.
Meridional overturning stream function in (a–c) the world ocean,
(d–f) Atlantic Ocean and (g–i) Indo-Pacific for PI, year 2050,
and for new steady-state conditions for a representative ensemble member reaching
the 1.5 ∘C warming target (columns). Circulation is clockwise along
positive (red) contours and anticlockwise along the negative (blue) contours.
Changes in temperature, salinity and density at AD 2100 (a–c),
at AD 3150 (d–f) and for new steady-state conditions (g–i)
compared to pre-industrial conditions.
Changes in O2 and its components for new steady-state conditions
relative to preindustrial steady state for a single representative ensemble
member reaching the 1.5 ∘C warming target. (a) Change in total
O2, (b) change in O2 due to biology, (c) change
in O2 due to solubility, (d) changes in ideal age.
The authors declare that they have no conflict of interest.
This article is part of the special issue “The Earth system at a
global warming of 1.5 ∘C and 2.0 ∘C”. It is not associated
with a conference.
Acknowledgements
We thank Andreas Oschlies and Andreas Schmittner for their thorough reviews
of our manuscript, which allowed us to better communicate our results. We
also thank Patrik Pfister and Thomas Frölicher for fruitful discussions.
This work was supported by the Swiss National Science Foundation (200020_172476).
Edited by: Axel Kleidon
Reviewed by: Andreas Oschlies and Andreas Schmittner
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