255
This chapter should be cited as:
Rhein, M., S.R. Rintoul, S. Aoki, E. Campos, D. Chambers, R.A. Feely, S. Gulev, G.C. Johnson, S.A. Josey, A. Kostianoy,
C. Mauritzen, D. Roemmich, L.D. Talley and F. Wang, 2013: Observations: Ocean. In: Climate Change 2013: The
Physical Science Basis. Contribution of Working Group I to the Fifth Assessment Report of the Intergovernmental
Panel on Climate Change [Stocker, T.F., D. Qin, G.-K. Plattner, M. Tignor, S.K. Allen, J. Boschung, A. Nauels, Y. Xia, V.
Bex and P.M. Midgley (eds.)]. Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA.
Coordinating Lead Authors:
Monika Rhein (Germany), Stephen R. Rintoul (Australia)
Lead Authors:
Shigeru Aoki (Japan), Edmo Campos (Brazil), Don Chambers (USA), Richard A. Feely (USA),
Sergey Gulev (Russian Federation), Gregory C. Johnson (USA), Simon A. Josey (UK), Andrey
Kostianoy (Russian Federation), Cecilie Mauritzen (Norway), Dean Roemmich (USA), Lynne D.
Talley (USA), Fan Wang (China)
Contributing Authors:
Ian Allison (Australia), Michio Aoyama (Japan), Molly Baringer (USA), Nicholas R. Bates
(Bermuda), Timothy Boyer (USA), Robert H. Byrne (USA), Sarah Cooley (USA), Stuart Cunningham
(UK), Thierry Delcroix (France), Catia M. Domingues (Australia), Scott Doney (USA), John Dore
(USA), Paul. J. Durack (USA/Australia), Rana Fine (USA), Melchor González-Dávila (Spain), Simon
Good (UK), Nicolas Gruber (Switzerland), Mark Hemer (Australia), David Hydes (UK), Masayoshi
Ishii (Japan), Stanley Jacobs (USA), Torsten Kanzow (Germany), David Karl (USA), Georg Kaser
(Austria/Italy), Alexander Kazmin (Russian Federation), Robert Key (USA), Samar Khatiwala
(USA), Joan Kleypas (USA), Ronald Kwok (USA), Kitack Lee (Republic of Korea), Eric Leuliette
(USA), Melisa Menéndez (Spain), Calvin Mordy (USA), Jon Olafsson (Iceland), James Orr (France),
Alejandro Orsi (USA), Geun-Ha Park (Republic of Korea), Igor Polyakov (USA), Sarah G. Purkey
(USA), Bo Qiu (USA), Gilles Reverdin (France), Anastasia Romanou (USA), Sunke Schmidtko
(UK), Raymond Schmitt (USA), Koji Shimada (Japan), Doug Smith (UK), Thomas M. Smith (USA),
Uwe Stöber (Germany), Lothar Stramma (Germany), Toshio Suga (Japan), Neil Swart (Canada/
South Africa), Taro Takahashi (USA), Toste Tanhua (Germany), Karina von Schuckmann (France),
Hans von Storch (Germany), Xiaolan Wang (Canada), Rik Wanninkhof (USA), Susan Wijffels
(Australia), Philip Woodworth (UK), Igor Yashayaev (Canada), Lisan Yu (USA)
Review Editors:
Howard Freeland (Canada), Silvia Garzoli (USA), Yukihiro Nojiri (Japan)
Observations: Ocean
3
256
3
Table of Contents
Executive Summary ..................................................................... 257
3.1 Introduction ...................................................................... 260
3.2 Changes in Ocean Temperature and
Heat Content .................................................................... 260
3.2.1 Effects of Sampling on Ocean Heat
Content Estimates ..................................................... 260
3.2.2 Upper Ocean Temperature......................................... 261
3.2.3 Upper Ocean Heat Content ....................................... 262
3.2.4 Deep Ocean Temperature and Heat Content ............. 263
3.2.5 Conclusions ............................................................... 263
Box 3.1: Change in Global Energy Inventory ......................... 264
3.3 Changes in Salinity and Freshwater Content ......... 265
3.3.1 Introduction .............................................................. 265
3.3.2 Global to Basin-Scale Trends ..................................... 267
3.3.3 Regional Changes in Upper Ocean Salinity ............... 271
3.3.4 Evidence for Change of the Hydrological Cycle
from Salinity Changes ............................................... 273
3.3.5 Conclusions ............................................................... 273
3.4 Changes in Ocean Surface Fluxes ............................... 273
3.4.1 Introduction .............................................................. 273
3.4.2 Air–Sea Heat Fluxes .................................................. 274
3.4.3 Ocean Precipitation and Freshwater Flux .................. 275
3.4.4 Wind Stress ............................................................... 276
3.4.5 Changes in Surface Waves ........................................ 277
3.4.6 Conclusions ............................................................... 278
3.5 Changes in Water-Mass Properties ............................ 278
3.5.1 Introduction .............................................................. 278
3.5.2 Intermediate Waters .................................................. 278
3.5.3 Deep and Bottom Waters .......................................... 279
3.5.4 Conclusions ............................................................... 280
3.6 Changes in Ocean Circulation ..................................... 281
3.6.1 Global Observations of Ocean
Circulation Variability ................................................ 281
3.6.2 Wind-Driven Circulation Variability in the
Pacific Ocean ............................................................. 281
3.6.3 The Atlantic Meridional Overturning Circulation ....... 282
3.6.4 The Antarctic Meridional Overturning Circulation ..... 284
3.6.5 Water Exchange Between Ocean Basins ................... 284
3.6.6 Conclusions ............................................................... 285
3.7 Sea Level Change, Including Extremes ..................... 285
3.7.1 Introduction and Overview of Sea Level
Measurements .......................................................... 285
3.7.2 Trends in Global Mean Sea Leve
and Components ....................................................... 286
3.7.3 Regional Distribution of Sea Level Change ............... 288
3.7.4 Assessment of Evidence for Accelerations in
Sea Level Rise ........................................................... 289
3.7.5 Changes in Extreme Sea Level .................................. 290
3.7.6 Conclusions ............................................................... 291
3.8 Ocean Biogeochemical Changes, Including
Anthropogenic Ocean Acidification ........................... 291
3.8.1 Carbon ...................................................................... 292
3.8.2 Anthropogenic Ocean Acidification ........................... 293
3.8.3 Oxygen ...................................................................... 294
Box 3.2: Ocean Acidification .................................................... 295
3.8.4 Nutrients ................................................................... 298
3.8.5 Conclusions ............................................................... 300
3.9 Synthesis ............................................................... 301
References .................................................................................. 303
Appendix 3.A: Availability of Observations for
Assessment of Change in the Oceans.................................... 311
3.A.1 Subsurface Ocean Temperature and Heat Content .... 311
3.A.2 Salinity ...................................................................... 312
3.A.3 Sea Level ................................................................... 312
3.A.4 Biogeochemistry ....................................................... 312
Frequently Asked Questions
FAQ 3.1 Is the Ocean Warming? .......................................... 266
FAQ 3.2 Is There Evidence for Changes in the
Earth’s Water Cycle? .............................................. 269
FAQ 3.3 How Does Anthropogenic Ocean Acidification
Relate to Climate Change? ................................... 297
257
3
Observations: Ocean Chapter 3
1
In this Report, the following terms have been used to indicate the assessed likelihood of an outcome or a result: Virtually certain 99–100% probability, Very likely 90–100%,
Likely 66–100%, About as likely as not 33–66%, Unlikely 0–33%, Very unlikely 0-10%, Exceptionally unlikely 0–1%. Additional terms (Extremely likely: 95–100%, More likely
than not >50–100%, and Extremely unlikely 0–5%) may also be used when appropriate. Assessed likelihood is typeset in italics, e.g., very likely (see Section 1.4 and Box TS.1
for more details).
2
In this Report, the following summary terms are used to describe the available evidence: limited, medium, or robust; and for the degree of agreement: low, medium, or high.
A level of confidence is expressed using five qualifiers: very low, low, medium, high, and very high, and typeset in italics, e.g., medium confidence. For a given evidence and
agreement statement, different confidence levels can be assigned, but increasing levels of evidence and degrees of agreement are correlated with increasing confidence (see
Section 1.4 and Box TS.1 for more details).
Executive Summary
Temperature and Heat Content Changes
It is virtually certain
1
that the upper ocean (above 700 m) has
warmed from 1971 to 2010, and likely that it has warmed from
the 1870s to 1971. Confidence in the assessment for the time period
since 1971 is high
2
based on increased data coverage after this date
and on a high level of agreement among independent observations of
subsurface temperature [3.2], sea surface temperature [2.4.2], and sea
level rise, which is known to include a substantial component due to
thermal expansion [3.7, Chapter 13]. There is less certainty in changes
prior to 1971 because of relatively sparse sampling in earlier time peri-
ods. The strongest warming is found near the sea surface (0.11 [0.09
to 0.13] °C per decade in the upper 75 m between 1971 and 2010),
decreasing to about 0.015°C per decade at 700 m. It is very likely that
the surface intensification of this warming signal increased the ther-
mal stratification of the upper ocean by about 4% between 0 and 200
m depth. Instrumental biases in historical upper ocean temperature
measurements have been identified and reduced since AR4, diminish-
ing artificial decadal variation in temperature and upper ocean heat
content, most prominent during the 1970s and 1980s. {3.2.1–3.2.3,
Figures 3.1, 3.2 and 3.9}
It is likely that the ocean warmed between 700 and 2000 m
from 1957 to 2009, based on 5-year averages. It is likely that
the ocean warmed from 3000 m to the bottom from 1992 to
2005, while no significant trends in global average tempera-
ture were observed between 2000 and 3000 m depth during
this period. Warming below 3000 m is largest in the Southern Ocean
{3.2.4, 3.5.1, Figures 3.2b and 3.3, FAQ 3.1}
It is virtually certain that upper ocean (0 to 700 m) heat content
increased during the relatively well-sampled 40-year period
from 1971 to 2010. Published rates for that time period range from
74 TW to 137 TW, with generally smaller trends for estimates that
assume zero anomalies in regions with sparse data. Using a statistical
analysis of ocean variability to estimate change in sparsely sampled
areas and to estimate uncertainties results in a rate of increase of
global upper ocean heat content of 137 [120–154] TW (medium confi-
dence). Although not all trends agree within their statistical uncertain-
ties, all are positive, and all are statistically different from zero. {3.2.3,
Figure 3.2}
Warming of the ocean between 700 and 2000 m likely contrib-
uted about 30% of the total increase in global ocean heat con-
tent (0 to 2000 m) between 1957 and 2009. Although globally
integrated ocean heat content in some of the 0 to 700 m estimates
increased more slowly from 2003 to 2010 than over the previous
decade, ocean heat uptake from 700 to 2000 m likely continued una-
bated during this period. {3.2.4, Figure 3.2, Box 9.2}
Ocean warming dominates the global energy change inventory.
Warming of the ocean accounts for about 93% of the increase
in the Earth’s energy inventory between 1971 and 2010 (high
confidence), with warming of the upper (0 to 700 m) ocean
accounting for about 64% of the total. Melting ice (including Arctic
sea ice, ice sheets and glaciers) and warming of the continents and
atmosphere account for the remainder of the change in energy. The
estimated net increase in the Earth’s energy storage between 1971 and
2010 is 274 [196 to 351] ZJ (1 ZJ = 10
21
Joules), with a heating rate of
213 TW from a linear fit to annual inventories over that time period,
equivalent to 0.42 W m
–2
heating applied continuously over the Earth’s
entire surface, and 0.55 W m
–2
for the portion due to ocean warming
applied over the ocean surface area. {Section 3.2.3, Figure 3.2, Box 3.1}
Salinity and Freshwater Content Changes
It is very likely that regional trends have enhanced the mean
geographical contrasts in sea surface salinity since the 1950s:
saline surface waters in the evaporation-dominated mid-
latitudes have become more saline, while relatively fresh sur-
face waters in rainfall-dominated tropical and polar regions
have become fresher. The mean contrast between high- and low-
salinity regions increased by 0.13 [0.08 to 0.17] from 1950 to 2008.
It is very likely that the interbasin contrast in freshwater content has
increased: the Atlantic has become saltier and the Pacific and Southern
oceans have freshened. Although similar conclusions were reached in
AR4, recent studies based on expanded data sets and new analysis
approaches provide high confidence in the assessment of trends in
ocean salinity. {3.3.2, 3.3.3, 3.3.5, Figures 3.4, 3.5 and 3.21d, FAQ 3.2}
It is very likely that large-scale trends in salinity have also
occurred in the ocean interior. It is likely that both the subduction
of surface water anomalies formed by changes in evaporation – pre-
cipitation (E – P) and the movement of density surfaces due to warm-
ing have contributed to the observed changes in subsurface salinity.
{3.3.2–3.3.4, Figures 3.5 and 3.9}
The spatial patterns of the salinity trends, mean salinity and
the mean distribution of E – P are all similar. This provides, with
medium confidence, indirect evidence that the pattern of E – P over the
oceans has been enhanced since the 1950s.{3.3.2–3.3.4, Figures 3.4,
3.5 and 3.20d, FAQ 3.2}.
258
Chapter 3 Observations: Ocean
3
Air–Sea Flux and Wave Height Changes
Uncertainties in air–sea heat flux data sets are too large to allow
detection of the change in global mean net air-sea heat flux, of
the order of 0.5 W m
–2
since 1971, required for consistency with
the observed ocean heat content increase. The products cannot
yet be reliably used to directly identify trends in the regional or global
distribution of evaporation or precipitation over the oceans on the time
scale of the observed salinity changes since 1950. {3.4.2, 3.4.3, Figures
3.6 and 3.7}
Basin-scale wind stress trends at decadal to centennial time
scales have been observed in the North Atlantic, Tropical Pacif-
ic and Southern Ocean with low to medium confidence. These
results are based largely on atmospheric reanalyses, in some cases a
single product, and the confidence level is dependent on region and
time scale considered. The evidence is strongest for the Southern
Ocean, for which there is medium confidence that zonal mean wind
stress has increased in strength since the early 1980s. {3.4.4, Figure
3.8}
There is medium confidence based on ship observations and
reanalysis forced wave model hindcasts that mean significant
wave height has increased since the 1950s over much of the
North Atlantic north of 45°N, with typical winter season trends
of up to 20 cm per decade. {3.4.5}
Changes in Water Masses and Circulation
Observed changes in water mass properties likely reflect the
combined effect of long-term trends in surface forcing (e.g.,
warming of the surface ocean and changes in E – P) and inter-
annual-to-multi-decadal variability related to climate modes.
Most of the observed temperature and salinity changes in the ocean
interior can be explained by subduction and spreading of water masses
with properties that have been modified at the sea surface. From 1950
to 2000, it is likely that subtropical salinity maximum waters became
more saline, while fresh intermediate waters formed at higher latitude
have generally become fresher. For Upper North Atlantic Deep Water
changes in properties and formation rates are very likely dominated
by decadal variability. The Lower North Atlantic Deep Water has likely
cooled from 1955 to 2005, and the freshening trend highlighted in AR4
reversed in the mid-1990s. It is likely that the Antarctic Bottom Water
warmed and contracted globally since the 1980s and freshened in the
Indian/Pacific sectors from 1970 to 2008. {3.5, FAQ 3.1}
Recent observations have strengthened evidence for variability
in major ocean circulation systems on time scales from years to
decades. It is very likely that the subtropical gyres in the North Pacific
and South Pacific have expanded and strengthened since 1993. It is
about as likely as not that this is linked to decadal variability in wind
forcing rather than being part of a longer-term trend. Based on meas-
urements of the full Atlantic Meridional Overturning Circulation and its
individual components at various latitudes and different time periods,
there is no evidence of a long-term trend. There is also no evidence for
trends in the transports of the Indonesian Throughflow, the Antarctic
Circumpolar Current (ACC), or between the Atlantic Ocean and Nordic
Seas. However, there is medium confidence that the ACC shifted south
between 1950 and 2010, at a rate equivalent to about 1° of latitude in
40 years. {3.6, Figures 3.10, 3.11}
Sea Level Change
Global mean sea level (GMSL) has risen by 0.19 [0.17 to 0.21] m
over the period 1901–2010, calculated using the mean rate over
these 110 years, based on tide gauge records and since 1993
additionally on satellite data. It is very likely that the mean
rate was 1.7 [1.5 to 1.9] mm yr
–1
between 1901 and 2010 and
increased to 3.2 [2.8 to 3.6] mm yr
–1
between 1993 and 2010.
This assessment is based on high agreement among multiple studies
using different methods, long tide gauge records corrected for verti-
cal land motion and independent observing systems (tide gauges and
altimetry) since 1993 (see also TFE.2, Figure 1). It is likely that GMSL
rose between 1920 and 1950 at a rate comparable to that observed
between 1993 and 2010, as individual tide gauges around the world
and reconstructions of GMSL show increased rates of sea level rise
during this period. Rates of sea level rise over broad regions can be
several times larger or smaller than that of GMSL for periods of sev-
eral decades due to fluctuations in ocean circulation. High agreement
between studies with and without corrections for vertical land motion
suggests that it is very unlikely that estimates of the global average
rate of sea level change are significantly biased owing to vertical land
motion that has been unaccounted for. {3.7.2, 3.7.3, Table 3.1, Figures
3.12, 3.13, 3.14}
It is very likely that warming of the upper 700 m has been con-
tributing an average of 0.6 [0.4 to 0.8] mm yr
–1
of sea level rise
since 1971. It is likely that warming between 700 m and 2000 m has
been contributing an additional 0.1 mm yr
–1
[0 to 0.2] of sea level rise
since 1971, and that warming below 2000 m has been contributing
another 0.1 [0.0 to 0.2] mm yr
–1
of sea level rise since the early 1990s.
{3.7.2, Figure 3.13}
It is likely that the rate of sea level rise increased from the early
19th century to the early 20th century, and increased further
over the 20th century. The inference of 19th century change is
based on a small number of very long tide gauge records from north-
ern Europe and North America. Multiple long tide gauge records and
reconstructions of global mean sea level confirm a higher rate of rise
from the late 19th century. It is likely that the average acceleration
over the 20th century is [–0.002 to 0.019] mm yr
–2
, as two of three
reconstructions extending back to at least 1900 show an acceleration
during the 20th century. {3.7.4}
It is likely that the magnitude of extreme high sea level events
has increased since 1970. A rise in mean sea level can explain most
of the increase in extreme sea levels: changes in extreme high sea
levels are reduced to less than 5 mm yr
–1
at 94% of tide gauges once
the rise in mean sea level is accounted for. {3.7.5, Figure 3.15}
259
3
Observations: Ocean Chapter 3
Changes in Ocean Biogeochemistry
Based on high agreement between independent estimates using
different methods and data sets (e.g., oceanic carbon, oxygen,
and transient tracer data), it is very likelythat the global ocean
inventory of anthropogenic carbon (C
ant
) increased from 1994
to 2010. The oceanic C
ant
inventory in 2010 is estimated to be 155
PgC with an uncertainty of ±20%. The annual global oceanic uptake
rates calculated from independent data sets (from oceanic C
ant
inven-
tory changes, from atmospheric O
2
/N
2
measurements or from partial
pressure of carbon dioxide (pCO
2
) data) and for different time periods
agree with each other within their uncertainties and very likely are in
the range of 1.0 to 3.2 PgC yr
–1
{3.8.1, Figure 3.16}
Uptake of anthropogenic CO
2
results in gradual acidification of
the ocean. The pH of surface seawater has decreased by 0.1 since
the beginning of the industrial era, corresponding to a 26% increase
in hydrogen ion concentration (high confidence). The observed pH
trends range between –0.0014 and –0.0024 yr
–1
in surface waters. In
the ocean interior, natural physical and biological processes, as well as
uptake of anthropogenic CO
2
, can cause changes in pH over decadal
and longer time scales. {3.8.2, Table 3.2, Box 3.2, Figures 3.18, 3.19,
FAQ 3.3}
High agreement among analyses provides medium confidence
that oxygen concentrations have decreased in the open ocean
thermocline in many ocean regions since the 1960s. The general
decline is consistent with the expectation that warming-induced strati-
fication leads to a decrease in the supply of oxygen to the thermocline
from near surface waters, that warmer waters can hold less oxygen,
and that changes in wind-driven circulation affect oxygen concentra-
tions. It is likely that the tropical oxygen minimum zones have expand-
ed in recent decades. {3.8.3, Figure 3.20}
Synthesis
The observations summarized in this chapter provide strong
evidence that ocean properties of relevance to climate have
changed during the past 40 years, including temperature, salin-
ity, sea level, carbon, pH, and oxygen. The observed patterns of
change in the subsurface ocean are consistent with changes in
the surface ocean in response to climate change and natural
variability and with known physical and biogeochemical pro-
cesses in the ocean, providing high confidence in this assess-
ment. {3.9, Figures 3.21, 3.22}
260
Chapter 3 Observations: Ocean
3
3.1 Introduction
The ocean influences climate by storing and transporting large
amounts of heat, freshwater, and carbon, and by exchanging these
properties with the atmosphere. About 93% of the excess heat energy
stored by the Earth over the last 50 years is found in the ocean (Church
et al., 2011; Levitus et al., 2012). The ability of the ocean to store vast
amounts of heat reflects the large mass and heat capacity of seawater
relative to air and the fact that ocean circulation connects the surface
and interior ocean. More than three quarters of the total exchange
of water between the atmosphere and the Earth’s surface through
evaporation and precipitation takes place over the oceans (Schmitt,
2008). The ocean contains 50 times more carbon than the atmosphere
(Sabine et al., 2004) and is at present acting to slow the rate of climate
change by absorbing about 30% of human emissions of carbon diox-
ide (CO
2
) from fossil fuel burning, cement production, deforestation
and other land use change (Mikaloff-Fletcher et al., 2006; Le Quéré et
al., 2010). Changes in the ocean may result in climate feedbacks that
either increase or reduce the rate of climate change. Climate variabil-
ity and change on time scales from seasons to millennia is therefore
closely linked to the ocean and its interactions with the atmosphere
and cryosphere. The large inertia of the oceans means that they nat-
urally integrate over short-term variability and often provide a clearer
signal of longer-term change than other components of the climate
system. Observations of ocean change therefore provide a means to
track the evolution of climate change, and a relevant benchmark for
climate models.
The lack of long-term measurements of the global ocean and chang-
es in the observing system over time makes documenting and under-
standing change in the oceans a difficult challenge (Appendix 3.A).
Many of the issues raised in Box 2.1 regarding uncertainty in atmos-
pheric climate records are common to oceanographic data. Despite the
limitations of historical records, AR4 identified significant trends in a
number of ocean variables relevant to climate change, including ocean
heat content, sea level, regional patterns of salinity, and biogeochem-
ical parameters (Bindoff et al., 2007). Since AR4, substantial progress
has been made in improving the quality and coverage of ocean obser-
vations. Biases in historical measurements have been identified and
reduced, providing a clearer record of past change. The Argo array of
profiling floats has provided near-global, year-round measurements of
temperature and salinity in the upper 2000 m since 2005. The satellite
altimetry record is now more than 20 years in length. Longer continu-
ous time series of important components of the meridional overturning
circulation and tropical oceans have been obtained. The spatial and
temporal coverage of biogeochemical measurements in the ocean has
expanded. As a result of these advances, there is now stronger evi-
dence of change in the ocean, and our understanding of the causes of
ocean change is improved.
This chapter summarizes the observational evidence of change in the
ocean, with an emphasis on basin- and global-scale changes relevant to
climate, with a focus on studies published since the AR4. As in Chapter
2, the robustness of observed changes is assessed relative to sources
of observational uncertainty. The attribution of ocean change, includ-
ing the degree to which observed changes are consistent with anthro-
pogenic climate change, is addressed in Chapter 10. The evidence for
changes in subsurface ocean temperature and heat content is assessed
in Section 3.2; changes in sea surface temperature (SST) are covered
in Chapter 2. Changes in ocean heat content dominate changes in the
global energy inventory (Box 3.1). Recent studies have strengthened
the evidence for regional changes in ocean salinity and their link to
changes in evaporation and precipitation over the oceans (Section 3.3),
a connection already identified in AR4. Evidence for changes in the
fluxes of heat, water and momentum (wind stress) across the air–sea
interface is assessed in Section 3.4. Considering ocean changes from
a water-mass perspective adds additional insight into the nature and
causes of ocean change (Section 3.5). Although direct observations
of ocean circulation are more limited than those of temperature and
salinity, there is growing evidence of variability and change of ocean
current patterns relevant to climate (Section 3.6). Observations of sea
level change are summarized in Section 3.7; Chapter 13 builds on the
evidence presented in this and other chapters to provide an overall
synthesis of past and future sea level change. Biogeochemical chang-
es in the ocean, including ocean acidification, are covered in Section
3.8. Chapter 6 combines observations with models to discuss past and
present changes in the carbon cycle. Section 3.9 provides an overall
synthesis of changes observed in the ocean during the instrumental
period and highlights key uncertainties. Unless otherwise noted, uncer-
tainties (in square brackets) represent 5 to 95% confidence intervals.
3.2 Changes in Ocean Temperature and
Heat Content
3.2.1 Effects of Sampling on Ocean Heat Content
Estimates
Temperature is the most often measured subsurface ocean variable.
Historically, a variety of instruments have been used to measure tem-
perature, with differing accuracies, precisions, and sampling depths.
Both the mix of instruments and the overall sampling patterns have
changed in time and space (Boyer et al., 2009), complicating efforts
to determine and interpret long-term change. The evolution of the
observing system for ocean temperature is summarized in Appendix
3.A. Upper ocean temperature (hence heat content) varies over multi-
ple time scales including seasonal (e.g., Roemmich and Gilson, 2009),
interannual (e.g. associated with El Niño, which has a strong influence
on ocean heat uptake, Roemmich and Gilson, 2011), decadal (e.g.,
Carson and Harrison, 2010), and centennial (Gouretski et al., 2012;
Roemmich et al., 2012). Ocean data assimilation products using these
data exhibit similar significant variations (e.g., Xue et al., 2012). Sparse
historical sampling coupled with large amplitude variations on shorter
time and spatial scales raise challenges for estimating globally aver-
aged upper ocean temperature changes. Uncertainty analyses indicate
that the historical data set begins to be reasonably well suited for this
purpose starting around 1970 (e.g., Domingues et al., 2008; Lyman and
Johnson, 2008; Palmer and Brohan, 2011). UOHC uncertainty estimates
shrink after 1970 with improved sampling, so this assessment focus-
es on changes since 1971. Estimates of UOHC have been extended
back to 1950 by averaging over longer time intervals, such as 5-year
running means, to compensate for sparse data distributions in earlier
time periods (e.g., Levitus et al., 2012). These estimates may be most
appropriate in the deeper ocean, where strong interannual variability
in upper ocean temperature distributions such as that associated with
El Niño (Roemmich and Gilson, 2011) is less likely to be aliased.
261
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Observations: Ocean Chapter 3
Since AR4 the significant impact of measurement biases in some of
the widely used instruments (the expendable (XBT) and mechanical
bathythermograph (MBT) as well as a subset of Argo floats) on esti-
mates of ocean temperature and upper (0 to 700 m) ocean heat con-
tent (hereafter UOHC) changes has been recognized (Gouretski and
Koltermann, 2007; Barker et al., 2011). Careful comparison of meas-
urements from the less accurate instruments with those from the more
accurate ones has allowed some of the biases to be identified and
reduced (Wijffels et al., 2008; Ishii and Kimoto, 2009; Levitus et al.,
2009; Gouretski and Reseghetti, 2010; Hamon et al., 2012). One major
consequence of this bias reduction has been the reduction of an artifi-
cial decadal variation in upper ocean heat content that was apparent
in the observational assessment for AR4, in notable contrast to climate
model output (Domingues et al., 2008). Substantial time-dependent
XBT and MBT biases introduced spurious warming in the 1970s and
cooling in the early 1980s in the analyses assessed in AR4. Most ocean
state estimates that assimilate biased data (Carton and Santorelli,
2008) also showed this artificial decadal variability while one (Stam-
mer et al., 2010) apparently rejected these data on dynamical grounds.
More recent estimates assimilating better-corrected data sets (Giese
et al., 2011) also result in reduced artificial decadal variability during
this time period.
Recent estimates of upper ocean temperature change also differ in
their treatment of unsampled regions. Some studies (e.g., Ishii and
Kimoto, 2009; Levitus et al., 2012) effectively assume a temperature
anomaly of zero in these regions, while other studies (Palmer et al.,
2007; Lyman and Johnson, 2008) assume that the averages of sampled
regions are representative of the global mean in any given year, and yet
others (Smith and Murphy, 2007; Domingues et al., 2008) use ocean
statistics (from numerical model output and satellite altimeter data,
respectively) to extrapolate temperature anomalies in sparsely sam-
pled areas and estimate uncertainties. These differences in approach,
coupled with choice of background climatology, can lead to significant
divergence in basin-scale averages (Gleckler et al., 2012), especially in
sparsely sampled regions (e.g., the extratropical Southern Hemisphere
(SH) prior to Argo), and as a result can produce different global averag-
es (Lyman et al., 2010). However, for well-sampled regions and times,
the various analyses of temperature changes yield results in closer
agreement, as do reanalyses (Xue et al., 2012).
3.2.2 Upper Ocean Temperature
Depth-averaged 0 to 700 m ocean temperature trends from 1971 to
2010 are positive over most of the globe (Levitus et al., 2009; Figure
3.1a). The warming is more prominent in the Northern Hemisphere
(NH), especially the North Atlantic. This result holds in different anal-
yses, using different time periods, bias corrections and data sources
(e.g., with or without XBT or MBT data) (e.g., Palmer et al., 2007;
Durack and Wijffels, 2010; Gleckler et al., 2012; Figures 3.1 and 3.9).
However, the greater volume of the SH oceans increases the contribu-
tion of their warming to global heat content. Zonally averaged upper
ocean temperature trends show warming at nearly all latitudes and
depths (Levitus et al., 2009, Figure 3.1b). A maximum in warming south
of 30°S appears in Figure 3.1b, but is not as strong as in other analyses
(e.g., Gille, 2008), likely because the data are relatively sparse in this
location so anomalies are attenuated by the objectively analyzed fields
Figure 3.1 | (a) Depth-averaged 0 to 700 m temperature trend for 1971–2010
(longitude vs. latitude, colours and grey contours in degrees Celsius per decade). (b)
Zonally averaged temperature trends (latitude vs. depth, colours and grey contours in
degrees Celsius per decade) for 1971–2010 with zonally averaged mean temperature
over-plotted (black contours in degrees Celsius). (c) Globally averaged temperature
anomaly (time vs. depth, colours and grey contours in degrees Celsius) relative to the
1971–2010 mean. (d) Globally averaged temperature difference between the ocean
surface and 200 m depth (black: annual values, red: 5-year running mean). All panels
are constructed from an update of the annual analysis of Levitus et al. (2009).
(a)
80°S 60°S40°S 20°S0°S 20°N 40°N 60°N 80°N
Latitude
Depth (m)
(b)
0
2
4
6
0
2
4
4
6
8
8
10
12
14
16
18
20
22
24
26
100
200
0
Depth (m)
(c)
1960 1970 1980 1990 2000 2010
700
600
500
400
300
200
100
0
6.1
6.3
6.5
6.7
Year
T0−T200 (°C)
(d)
(a,b) Temp. trend (°C per decade)
(c) Temp. anom. (°C)
−0.3
−0.25
−0.2
−0.15
−0.1
−0.05
0
0.05
0.1
0.15
0.2
0.25
0.3
300
400
500
600
700
used for Figure 3.1 and because warming in the upper 1000 m of the
Southern Ocean was stronger between the 1930s and the 1970s than
between the 1970s and 1990s (Gille, 2008). Another warming maxi-
mum is present at 25°N to 65°N. Both warming signals extend to 700
m (Levitus et al., 2009, Figure 3.1b), and are consistent with poleward
displacement of the mean temperature field. Other zonally averaged
temperature changes are also consistent with poleward displacement
of the mean temperatures. For example, cooling at depth between
30°S and the equator (Figure 3.1b) is consistent with a southward shift
of cooler water near the equator. Poleward displacements of some sub-
tropical and subpolar zonal currents and associated wind changes are
discussed in Section 3.6.
Globally averaged ocean temperature anomalies as a function of depth
and time (Figure 3.1c) relative to a 1971–2010 mean reveal warm-
ing at all depths in the upper 700 m over the relatively well-sampled
40-year period considered. Strongest warming is found closest to the
262
Chapter 3 Observations: Ocean
3
sea surface, and the near-surface trends are consistent with indepen-
dently measured SST (Chapter 2). The global average warming over
this period is 0.11 [0.09 to 0.13] °C per decade in the upper 75 m,
decreasing to 0.015°C per decade by 700 m (Figure 3.1c). Comparison
of Argo data to Challenger expedition data from the 1870s suggests
that warming started earlier than 1971, and was also larger in the
Atlantic than in the Pacific over that longer time interval (Roemmich et
al., 2012). An observational analysis of temperature in the upper 400
m of the global ocean starting in the year 1900 (Gouretski et al., 2012)
finds warming between about 1900 and 1945, as well as after 1970,
with some evidence of slight cooling between 1945 and 1970.
The globally averaged temperature difference between the ocean sur-
face and 200 m (Figure 3.1d) increased by about 0.25ºC from 1971 to
2010 (Levitus et al., 2009). This change, which corresponds to a 4%
increase in density stratification, is widespread in all the oceans north
of about 40°S.
A potentially important impact of ocean warming is the effect on sea
ice, floating glacial ice and ice sheet dynamics (see Chapter 4 for a
discussion of these topics). Although some of the global integrals of
UOHC neglect changes poleward of ±60° (Ishii and Kimoto, 2009) or
±65° (Domingues et al., 2008) latitude, at least some parts of the Arctic
have warmed: In the Arctic Ocean, subsurface pulses of relatively warm
water of Atlantic origin can be traced around the Eurasian Basin, and
analyses of data from 1950–2010 show a decadal warming of this
water mass since the late 1970s (Polyakov et al., 2012), as well as a
shoaling, by 75 to 90 m (Polyakov et al., 2010). Arctic surface waters
have also warmed, at least in the Canada Basin, from 1993 to 2007
(Jackson et al., 2010).
3.2.3 Upper Ocean Heat Content
Global integrals of 0 to 700 m UOHC (Figure 3.2a) estimated from ocean
temperature measurements all show a gain from 1971 to 2010 (Palmer
et al., 2007; Smith and Murphy, 2007; e.g., Domingues et al., 2008; Ishii
and Kimoto, 2009; Levitus et al., 2012) . These estimates usually start
around 1950, although as noted in Section 3.2.1 and discussed in the
Appendix, historical data coverage is sparse, so global integrals are
increasingly uncertain for earlier years, especially prior to 1970. There
is some convergence towards agreement in instrument bias correction
algorithms since AR4 (Section 3.2.1), but other sources of uncertainty
include the different assumptions regarding mapping and integrating
UOHCs in sparsely sampled regions, differences in quality control of
temperature data, and differences among baseline climatologies used
for estimating changes in heat content (Lyman et al., 2010). Although
there are still apparent interannual variations about the upward trend
of global UOHC since 1970, different global estimates have variations
at different times and for different periods, suggesting that sub-decadal
variability in the time rate of change is still quite uncertain in the his-
torical record. Most of the estimates in Figure 3.2a do exhibit decreas-
es for a few years immediately following major volcanic eruptions in
1963, 1982 and 1991 (Domingues et al., 2008).
Again, all of the global integrals of UOHC in Figure 3.2a have increased
between 1971 and 2010. Linear trends fit to the UOHC estimates for
the relatively well-sampled 40-year period from 1971 to 2010 estimate
Figure 3.2: | (a) Observation-based estimates of annual global mean upper (0 to 700
m) ocean heat content in ZJ (1 ZJ = 10
21
Joules) updated from (see legend): Levitus et
al. (2012), Ishii and Kimoto (2009), Domingues et al. (2008), Palmer et al. (2007) and
Smith and Murphy (2007). Uncertainties are shaded and plotted as published (at the
one standard error level, except one standard deviation for Levitus, with no uncertain-
ties provided for Smith). Estimates are shifted to align for 2006–2010, 5 years that are
well measured by Argo, and then plotted relative to the resulting mean of all curves
for 1971, the starting year for trend calculations. (b) Observation-based estimates of
annual 5-year running mean global mean mid-depth (700 to 2000 m) ocean heat con-
tent in ZJ (Levitus et al., 2012) and the deep (2000 to 6000 m) global ocean heat
content trend from 1992 to 2005 (Purkey and Johnson, 2010), both with one standard
error uncertainties shaded (see legend).
1950 1960 1970 1980 1990 2000 2010
−100
−50
0
50
100
150
Year
0−700 m OHC (ZJ)
(a)
Levitus
Ishii
Domingues
Palmer
Smith
1950 1960 1970 1980 1990 2000 2010
−50
0
50
Deep OHC (ZJ)
(b)
700−2000 m
2000−6000 m
the heating rate required to account for this warming: 118 [82 to 154]
TW (1 TW = 10
12
watts) for Levitus et al. (2012), 98 [67 to 130] TW
for Ishii and Kimoto (2009), 137 [120 to 154] TW for Domingues et
al. (2008), 108 [80 to 136] TW for Palmer et al. (2007), and 74 [43 to
105] TW for Smith and Murphy (2007). Uncertainties are calculated
as 90% confidence intervals for an ordinary least squares fit, taking
into account the reduction in the degrees of freedom implied by the
temporal correlation of the residuals. Although these rates of energy
gain do not all agree within their statistical uncertainties, all are pos-
itive, and all are statistically different from zero. Generally the smaller
trends are for estimates that assume zero anomalies in areas of sparse
data, as expected for that choice, which will tend to reduce trends and
variability. Hence the assessment of the Earth’s energy uptake (Box
3.1) employs a global UOHC estimate (Domingues et al., 2008) chosen
because it fills in sparsely sampled areas and estimates uncertainties
using a statistical analysis of ocean variability patterns.
Globally integrated ocean heat content in three of the five 0 to 700 m
estimates appear to be increasing more slowly from 2003 to 2010 than
over the previous decade (Figure 3.2a). Although this apparent change
263
3
Observations: Ocean Chapter 3
is concurrent with a slowing of the increase global mean surface tem-
perature, as discussed in Box 9.2, this is also a time period when the
ocean observing system transitioned from predominantly XBT to pre-
dominantly Argo temperature measurements (Johnson and Wijffels,
2011). Shifts in observing systems can sometimes introduce spurious
signals, so this apparent recent change should be viewed with caution.
3.2.4 Deep Ocean Temperature and Heat Content
Below 700 m data coverage is too sparse to produce annual global
ocean heat content estimates prior to about 2005, but from 2005 to
2010 and 0 to 1500 m the global ocean warmed (von Schuckmann and
Le Traon, 2011). Five-year running mean estimates yield a 700 to 2000
m global ocean heat content trend from 1957 to 2009 (Figure 3.2b)
that is about 30% of that for 0 to 2000 m over the length of the record
(Levitus et al., 2012). Ocean heat uptake from 700 to 2000 m likely
continues unabated since 2003 (Figure 3.2b); as a result, ocean heat
content from 0 to 2000 m shows less slowing after 2003 than does 0
to 700 m heat content (Levitus et al., 2012).
Global sampling of the ocean below 2000 m is limited to a number
of repeat oceanographic transects, many occupied only in the last
few decades (Figure 3.3b), and several time-series stations, some of
which extend over decades. This sparse sampling in space and time
makes assessment of global deep ocean heat content variability less
certain than that for the upper ocean (Ponte, 2012), especially at mid-
depths, where vertical gradients are still sufficiently large for transient
variations (ocean eddies, internal waves, and internal tides) to alias
estimates made from sparse data sets. However, the deep North Atlan-
tic Ocean is better sampled than the rest of the globe, making esti-
mates of full-depth deep ocean heat content changes there feasible
north of 20ºN since the mid-1950s (Mauritzen et al., 2012).
Based on the limited information available, it is likely that the global
ocean did not show a significant temperature trend between 2000 and
3000 m depth from about 1992–2005 (Figures 3.2b and 3.3a; Kouketsu
et al., 2011). At these depths it has been around a millennium on aver-
age since waters in the Indian and Pacific Oceans were last exposed to
air–sea interaction (Gebbie and Huybers, 2012).
Warming from 1992 to 2005 is likely greater than zero from 3000 m
to the ocean floor (Figures 3.2b and 3.3a; Kouketsu et al., 2011), espe-
cially in recently formed Antarctic Bottom Water (AABW). South of the
Sub-Antarctic Front (Figure 3.3b), much of the water column warmed
between 1992 and 2005 (Purkey and Johnson, 2010). Globally, deep
warming rates are highest near 4500 m (Figure 3.3a), usually near
the sea floor where the AABW influence is strongest, and attenuate
towards the north (Figure 3.3b), where the AABW influence weakens.
Global scale abyssal warming on relatively short multi-decadal time
scales is possible because of communication of signals by planetary
waves originating within the Southern Ocean, reaching even such
remote regions as the North Pacific (Kawano et al., 2010; Masuda et
al., 2010). This AABW warming may partly reflect a recovery from cool
conditions induced by the 1970s Weddell Sea Polynya (Robertson et
al., 2002), but further north, in the Vema Channel of the South Atlantic,
observations since 1970 suggest strong bottom water warming did not
commence there until about 1991 (Zenk and Morozov, 2007).
Figure 3.3 | (a) Areal mean warming rates (ºC per decade) versus depth (thick lines)
with 5 to 95% confidence limits (shading), both global (orange) and south of the
Sub-Antarctic Front (purple), centred on 1992–2005. (b) Mean warming rates (ºC per
decade) below 4000 m (colour bar) estimated for deep ocean basins (thin black out-
lines), centred on 1992–2005. Stippled basin warming rates are not significantly differ-
ent from zero at 95% confidence. The positions of the Sub-Antarctic Front (purple line)
and the repeat oceanographic transects from which these warming rates are estimated
(thick black lines) also shown. (Data from Purkey and Johnson, 2010.)
−0.02 −0.01 0 0.01 0.02 0.03 0.04 0.05 0.06 0.07
1000
2000
3000
4000
5000
6000
Warming rate (°C per decade)
Depth (m)
(a)
Global
Southern
(b)
(°C per decade)
−0.05 0 0.05
In the North Atlantic, strong decadal variability in North Atlantic Deep
Water (NADW) temperature and salinity (Wang et al., 2010), largely
associated with the North Atlantic Oscillation (NAO, Box 2.5) (e.g., Yas-
hayaev, 2007; Sarafanov et al., 2008), complicates efforts to determine
long-term trends from the historical record. Heat content in the North
Atlantic north of 20°N from 2000 m to the ocean floor increased slight-
ly from 1955 to 1975, and then decreased more strongly from 1975
to 2005 (Mauritzen et al., 2012), with a net cooling trend of –4 TW
from 1955–2005 estimated from a linear fit. The global trend estimate
below 2000 m is +35 TW from 1992 to 2005 (Purkey and Johnson,
2010), with strong warming in the Southern Ocean.
3.2.5 Conclusions
It is virtually certain that the upper ocean (0 to 700 m) warmed from
1971 to 2010. This result is supported by three independent and con-
sistent methods of observation including (1) multiple analyses of
subsurface temperature measurements described here; (2) SST data
(Section 2.4.2) from satellites and in situ measurements from surface
drifters and ships; and (3) the record of sea level rise, which is known
to include a substantial component owing to thermosteric expansion
(Section 3.7 and Chapter 13). The warming rate is 0.11 [0.09 to 0.13]°C
per decade in the upper 75 m, decreasing to about 0.015°C per decade
by 700 m. It is very likely that surface intensification of the warming
264
Chapter 3 Observations: Ocean
3
Box 3.1 | Change in Global Energy Inventory
The Earth has been in radiative imbalance, with less energy exiting the top of the atmosphere than entering, since at least about 1970
(Murphy et al., 2009; Church et al., 2011; Levitus et al., 2012). Quantifying this energy gain is essential for understanding the response
of the climate system to radiative forcing. Small amounts of this excess energy warm the atmosphere and continents, evaporate water
and melt ice, but the bulk of it warms the ocean (Box 3.1, Figure 1). The ocean dominates the change in energy because of its large
mass and high heat capacity compared to the atmosphere. In addition, the ocean has a very low albedo and absorbs solar radiation
much more readily than ice.
The global atmospheric energy change inventory accounting for specific heating and water evaporation is estimated by combining
satellite estimates for temperature anomalies in the lower troposphere (Mears and Wentz, 2009a; updated to version 3.3) from 70°S
to 82.5°N and the lower stratosphere (Mears and Wentz, 2009b; updated to version 3.3) from 82.5°S to 82.5°N weighted by the ratio
of the portions of atmospheric mass they sample (0.87 and 0.13, respectively). These temperature anomalies are converted to energy
changes using a total atmospheric mass of 5.14 × 10
18
kg, a mean total water vapor mass of 12.7 × 10
15
kg (Trenberth and Smith, 2005),
a heat capacity of 1 J g
–1
°C
–1
, a latent heat of vaporization of 2.464
J kg
–1
and a fractional increase of integrated water vapor con-
tent of 0.075 °C
–1
(Held and Soden, 2006). Smaller changes in
potential and kinetic energy are considered negligible. Standard
deviations for each year of data are used for uncertainties, and
the time series starts in 1979. The warming trend from a linear fit
from 1979 to 2010 amounts to 2 TW (1 TW = 10
12
watts).
The global average rate of continental warming and its uncer-
tainty has been estimated from borehole temperature profiles
from 1500 to 2000 at 50-year intervals (Beltrami et al., 2002).
The 1950–2000 estimate of land warming, 6 TW, is extended into
the first decade of the 21st century, although that extrapolation
is almost certainly an underestimate of the energy absorbed, as
land surface air temperatures for years since 2000 are some of
the warmest on record (Section 2.4.1).
All annual ice melt rates (for glaciers and ice-caps, ice sheets
and sea ice from Chapter 4) are converted into energy change
using a heat of fusion (334 × 10
3
J kg
–1
) and density (920 kg
m
–3
) for freshwater ice. The heat of fusion and density of ice may
vary, but only slightly among the different ice types, and warm-
ing the ice from sub-freezing temperatures requires much less
energy than that to melt it, so these second-order contributions
are neglected here. The linear trend of energy storage from 1971
to 2010 is 7 TW.
For the oceans, an estimate of global upper (0 to 700 m depth)
ocean heat content change using ocean statistics to extrapo-
late to sparsely sampled regions and estimate uncertainties
(Domingues et al., 2008) is used (see Section 3.2), with a linear
trend from 1971 to 2010 of 137 TW. For the ocean from 700 to
2000 m, annual 5-year running mean estimates are used from
1970 to 2009 and annual estimates for 2010–2011 (Levitus et
al., 2012). For the ocean from 2000 m to bottom, a uniform rate
of energy gain of 35 [6 to 61] TW from warming rates centred on
1992–2005 (Purkey and Johnson, 2010) is applied from 1992 to
2011, with no warming below 2000 m assumed prior to 1992.
Their 5 to 95% uncertainty estimate may be too small, as it
(continued on next page)
Box 3.1, Figure 1 | Plot of energy accumulation in ZJ (1 ZJ = 10
21
J) within
distinct components of the Earth’s climate system relative to 1971 and from 1971
to 2010 unless otherwise indicated. See text for data sources. Ocean warming
(heat content change) dominates, with the upper ocean (light blue, above 700 m)
contributing more than the mid-depth and deep ocean (dark blue, below 700 m;
including below 2000 m estimates starting from 1992). Ice melt (light grey; for
glaciers and ice caps, Greenland and Antarctic ice sheet estimates starting from
1992, and Arctic sea ice estimate from 1979 to 2008); continental (land) warming
(orange); and atmospheric warming (purple; estimate starting from 1979) make
smaller contributions. Uncertainty in the ocean estimate also dominates the total
uncertainty (dot-dashed lines about the error from all five components at 90%
confidence intervals).
1980 1990 2000 2010
−100
−50
0
50
100
150
200
250
300
Year
Energy (ZJ)
Upper ocean
Deep ocean
Ice
Land
Atmosphere
Uncertainty
265
3
Observations: Ocean Chapter 3
assumes the usually sparse sampling in each deep ocean basin analysed is representative of the mean trend in that basin. The linear
trend for heating the ocean below 700 m is 62 TW for 1971–2010.
It is virtually certain that the Earth has gained substantial energy from 1971 to 2010 — the estimated increase in energy inventory
between 1971 and 2010 is 274 [196 to 351] ZJ (1 ZJ = 10
21
J), with a rate of 213 TW from a linear fit to the annual values over that
time period (Box 3.1, Figure 1). An energy gain of 274 ZJ is equivalent to a heating rate of 0.42 W m
-2
applied continuously over the
surface area of the earth (5.10 × 10
14
m
2
). Ocean warming dominates the total energy change inventory, accounting for roughly 93% on
average from 1971 to 2010 (high confidence). The upper ocean (0-700 m) accounts for about 64% of the total energy change inventory.
Melting ice (including Arctic sea ice, ice sheets and glaciers) accounts for 3% of the total, and warming of the continents 3%. Warming
of the atmosphere makes up the remaining 1%. The 1971–2010 estimated rate of oceanic energy gain is 199 TW from a linear fit to
data over that time period, implying a mean heat flux of 0.55 W m
–2
across the global ocean surface area (3.60 × 10
14
m
2
). The Earth’s
net estimated energy increase from 1993 to 2010 is 163 [127 to 201] ZJ with a trend estimate of 275 TW. The ocean portion of the trend
for 1993–2010 is 257 TW, equivalent to a mean heat flux into the ocean of 0.71 W m
–2
over the global ocean surface area.
Box 3.1 (continued)
signal increased the thermal stratification of the upper ocean by about
4% (between 0 and 200 m depth) from 1971 to 2010. It is also likely
that the upper ocean warmed over the first half of the 20th century,
based again on these same three independent and consistent, although
much sparser, observations. Deeper in the ocean, it is likely that the
waters from 700 to 2000 m have warmed on average between 1957
and 2009 and likely that no significant trend was observed between
2000 and 3000 m from 1992 to 2005. It is very likely that the deep
(2000 m to bottom) North Atlantic Ocean north of 20°N warmed from
1955 to 1975, and then cooled from 1975 to 2005, with an overall
cooling trend. It is likely that most of the water column south of the
Sub-Antarctic Front warmed at a rate of about 0.03°C per decade from
1992 to 2005, and waters of Antarctic origin warmed below 3000 m at
a global average rate approaching 0.01°C per decade at 4500 m over
the same time period. For the deep ocean. Sparse sampling is the larg-
est source of uncertainty below 2000 m depth.
3.3 Changes in Salinity and Freshwater Content
3.3.1 Introduction
The ocean plays a pivotal role in the global water cycle: about 85% of
the evaporation and 77% of the precipitation occurs over the ocean
(Schmitt, 2008). The horizontal salinity distribution of the upper ocean
largely reflects this exchange of freshwater, with high surface salinity
generally found in regions where evaporation exceeds precipitation,
and low salinity found in regions of excess precipitation and runoff
(Figure 3.4a,b). Ocean circulation also affects the regional distribution
of surface salinity. The subduction (Section 3.5) of surface waters trans-
fers the surface salinity signal into the ocean interior, so that subsurface
salinity distributions are also linked to patterns of evaporation, precip-
itation and continental run-off at the sea surface. Melting and freezing
of ice (both sea ice and glacial ice) also influence ocean salinity.
Regional patterns and amplitudes of atmospheric moisture transport
could change in a warmer climate, because warm air can contain more
moisture (FAQ 3.2). The water vapour content of the troposphere likely
has increased since the 1970s, at a rate consistent with the observed
warming (Sections 2.4.4, 2.5.5 and 2.5.6).
It has not been possible to detect robust trends in regional precipita-
tion and evaporation over the ocean because observations over the
ocean are sparse and uncertain (Section 3.4.2). Ocean salinity, on the
other hand, naturally integrates the small difference between these
two terms and has the potential to act as a rain gauge for precip-
itation minus evaporation over the ocean (e.g., Lewis and Fofonoff,
1979; Schmitt, 2008; Yu, 2011; Pierce et al., 2012; Terray et al., 2012;
Section 10.4). Diagnosis and understanding of ocean salinity trends
is also important because salinity changes, like temperature changes,
affect circulation and stratification, and therefore the ocean’s capacity
to store heat and carbon as well as to change biological productivity.
Salinity changes also contribute to regional sea level change (Steele
and Ermold, 2007).
In AR4, surface and subsurface salinity changes consistent with a
warmer climate were highlighted, based on linear trends for the period
between 1955 and 1998 in the historical global salinity data set (Boyer
et al., 2005) as well as on more regional studies. In the early few dec-
ades the salinity data distribution was good in the NH, especially the
North Atlantic, but the coverage was poor in some regions such as the
central South Pacific, central Indian and polar oceans (Appendix 3.A).
However, Argo provides much more even spatial and temporal cover-
age in the 2000s. These additional observations, improvements in the
availability and quality of historical data and new analysis approaches
now allow a more complete assessment of changes in salinity.
‘Salinity’ refers to the weight of dissolved salts in a kilogram of sea-
water. Because the total amount of salt in the ocean does not change,
the salinity of seawater can be changed only by addition or removal of
fresh water. All salinity values quoted in the chapter are expressed on
the Practical Salinity Scale 1978 (PSS78) (Lewis and Fofonoff, 1979).
266
Chapter 3 Observations: Ocean
3
Frequently Asked Questions
FAQ 3.1 | Is the Ocean Warming?
Yes, the ocean is warming over many regions, depth ranges and time periods, although neither everywhere nor
constantly. The signature of warming emerges most clearly when considering global, or even ocean basin, averages
over time spans of a decade or more.
Ocean temperature at any given location can vary greatly with the seasons. It can also fluctuate substantially
from year to year—or even decade to decade—because of variations in ocean currents and the exchange of heat
between ocean and atmosphere.
Ocean temperatures have been recorded for centuries, but it was not until around 1971 that measurements were
sufficiently comprehensive to estimate the average global temperature of the upper several hundred meters of
the ocean confidently for any given year. In fact, before the international Argo temperature/salinity profiling float
array first achieved worldwide coverage in 2005, the global average upper ocean temperature for any given year
was sensitive to the methodology used to estimate it.
Global mean upper ocean temperatures have increased over decadal time scales from 1971 to 2010. Despite large
uncertainty in most yearly means, this warming is a robust result. In the upper 75 m of the ocean, the global average
warming trend has been 0.11 [0.09 to 0.13]°C per decade over this time. That trend generally lessens from the
surface to mid-depth, reducing to about 0.04°C per decade by 200 m, and to less than 0.02°C per decade by 500 m.
Temperature anomalies enter the subsurface ocean by paths in addition to mixing from above (FAQ3.1, Figure
1). Colder—hence denser—waters from high latitudes can sink from the surface, then spread toward the equator
beneath warmer, lighter, waters at lower latitudes. At a few locations—the northern North Atlantic Ocean and the
Southern Ocean around Antarctica—ocean water is cooled so much that it sinks to great depths, even to the sea
floor. This water then spreads out to fill much of the rest of the deep ocean. As ocean surface waters warm, these
sinking waters also warm with time, increasing temperatures in the ocean interior much more quickly than would
downward mixing of surface heating alone.
In the North Atlantic, the temperature of these deep waters varies from decade to decade—sometimes warming,
sometimes cooling—depending on prevailing winter atmospheric patterns. Around Antarctica, bottom waters have
warmed detectably from about 1992–2005, perhaps due to the strengthening and southward shift of westerly
winds around the Southern Ocean over the last several decades. This warming signal in the deepest coldest bottom
waters of the world ocean is detectable, although it weakens northward in the Indian, Atlantic and Pacific Oceans.
Deep warming rates are generally less pronounced than ocean surface rates (around 0.03ºC per decade since the
1990s in the deep and bottom waters around Antarctica, and smaller in many other locations). However, they occur
over a large volume, so deep ocean warming contributes significantly to the total increase in ocean heat.
Estimates of historical changes in global average ocean temperature have become more accurate over the past
several years, largely thanks to the recognition, and reduction, of systematic measurement errors. By carefully
comparing less accurate measurements with sparser, more accurate ones at adjacent locations and similar times,
scientists have reduced some spurious instrumental biases in the historical record. These improvements revealed
that the global average ocean temperature has increased much more steadily from year to year than was reported
prior to 2008. Nevertheless, the global average warming rate may not be uniform in time. In some years, the ocean
appears to warm faster than average; in others, the warming rate seems to slow.
The ocean’s large mass and high heat capacity allow it to store huge amounts of energy—more than 1000 times
that in the atmosphere for an equivalent increase in temperature. The Earth is absorbing more heat than it is
emitting back into space, and nearly all this excess heat is entering the oceans and being stored there. The ocean
has absorbed about 93% of the combined heat stored by warmed air, sea, and land, and melted ice between 1971
and 2010.
The ocean’s huge heat capacity and slow circulation lend it significant thermal inertia. It takes about a decade
for near-surface ocean temperatures to adjust in response to climate forcing (Section 12.5), such as changes in
greenhouse gas concentrations. Thus, if greenhouse gas concentrations could be held at present levels into the
future, increases in the Earth’s surface temperature would begin to slow within about a decade. However, deep
ocean temperature would continue to warm for centuries to millennia (Section 12.5), and thus sea levels would
continue to rise for centuries to millennia as well (Section 13.5). (continued on next page)
267
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Observations: Ocean Chapter 3
FAQ 3.1 (continued)
FAQ 3.1, Figure 1 | Ocean heat uptake pathways. The ocean is stratified, with the coldest, densest water in the deep ocean (upper panels: use map at top for orienta-
tion). Cold Antarctic Bottom Water (dark blue) sinks around Antarctica then spreads northward along the ocean floor into the central Pacific (upper left panel: red arrows
fading to white indicate stronger warming of the bottom water most recently in contact with the ocean surface) and western Atlantic oceans (upper right panel), as well
as the Indian Ocean (not shown). Less cold, hence lighter, North Atlantic Deep Water (lighter blue) sinks in the northern North Atlantic Ocean (upper right panel: red
and blue arrow in the deep water indicates decadal warming and cooling), then spreads south above the Antarctic Bottom Water. Similarly, in the upper ocean (lower
left panel shows Pacific Ocean detail, lower right panel the Atlantic), cool Intermediate Waters (cyan) sink in sub-polar regions (red arrows fading to white indicating
warming with time), before spreading toward the equator under warmer Subtropical Waters (green), which in turn sink (red arrows fading to white indicate stronger
warming of the intermediate and subtropical waters most recently in contact with the surface) and spread toward the equator under tropical waters, the warmest and
lightest (orange) in all three oceans. Excess heat or cold entering at the ocean surface (top curvy red arrows) also mixes slowly downward (sub-surface wavy red arrows).
N
S
S
N
A
BC
D
Equator
Surface
500m
1000m
Surface
500m
1000m
Antarctica Antarctica
Arctic
Arctic
Equator
Equator
2500m
5000m
Surface
2500m
5000m
Surface
West Atlantic Ocean
ACBD
Pacific Ocean
A
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t
a
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B
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W
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3.3.2 Global to Basin-Scale Trends
The salinity of near-surface waters is changing on global and basin
scales, with an increase in the more evaporative regions and a decrease
in the precipitation-dominant regions in almost all ocean basins.
3.3.2.1 Sea Surface Salinity
Multi-decadal trends in sea surface salinity have been documented in
studies published since AR4 (Boyer et al., 2007; Hosoda et al., 2009;
Roemmich and Gilson, 2009; Durack and Wijffels, 2010), confirm-
ing the trends reported in AR4 based mainly on Boyer et al. (2005).
The spatial pattern of surface salinity change is similar to the distri-
bution of surface salinity itself: salinity tends to increase in regions
of high mean salinity, where evaporation exceeds precipitation, and
tends to decrease in regions of low mean salinity, where precipitation
dominates (Figure 3.4). For example, salinity generally increased in the
surface salinity maxima formed in the evaporation-dominated subtrop-
ical gyres. The surface salinity minima at subpolar latitudes and the
intertropical convergence zones have generally freshened. Interbasin
salinity differences are also enhanced: the relatively salty Atlantic has
become more saline on average, while the relatively fresh Pacific has
become fresher (Figures 3.5 and 3.9). No well-defined trend is found
in the subpolar North Atlantic , which is dominated by decadal varia-
bility from atmospheric modes like the North Atlantic Oscillation (NAO,
Box 2.5). The 50-year salinity trends in Figure 3.4c, both positive and
negative, are statistically significant at the 99% level over 43.8% of
the global ocean surface (Durack and Wijffels, 2010); trends were less
significant over the remainder of the surface. The patterns of salinity
change in the complementary Hosoda et al. (2009) study of differences
between the periods 1960–1989 and 2003–2007 (Figure 3.4d), using a
different methodology, have a point-to-point correlation of 0.64 with
268
Chapter 3 Observations: Ocean
3
the Durack and Wijffels (2010) results, with significant differences only
in limited locations such as adjacent to the West Indies, Labrador Sea,
and some coastlines (Figure 3.4c and d).
It is very likely that the globally averaged contrast between regions of
high and low salinity relative to the global mean salinity has increased.
The contrast between high and low salinity regions, averaged over the
ocean area south of 70°N, increased by 0.13 [0.08 to 0.17] PSS78 from
1950 to 2008 using the data set of Durack and Wijffels (2010) , and by
0.12 [0.10 to 0.15] PSS78 using the data set of Boyer et al. (2009) with
the range reported in brackets signifying a 99% confidence interval
(Figure 3.21d).
3.3.2.2 Upper Ocean Subsurface Salinity
Compatible with observed changes in surface salinity, robust mul-
ti-decadal trends in subsurface salinity have been detected (Boyer et
al., 2005; Boyer et al., 2007; Steele and Ermold, 2007; Böning et al.,
2008; Durack and Wijffels, 2010; Helm et al., 2010; Wang et al., 2010).
Global, zonally averaged multi-decadal salinity trends (1950–2008) in
the upper 500 m (Figures 3.4, 3.5, 3.9 and Section 3.5) show salinity
increases at the salinity maxima of the subtropical gyres, freshening
of the low-salinity intermediate waters sinking in the Southern Ocean
(Subantarctic Mode Water and Antarctic Intermediate Water) and
North Pacific (North Pacific Intermediate Water). On average, the Pacific
freshened, and the Atlantic became more saline. These trends, shown
in Figures 3.5 and 3.9, are significant at a 95% confidence interval.
Freshwater content in the upper 500 m very likely changed, based
on the World Ocean Database 2009 (Boyer et al., 2009), analyzed by
Durack and Wijffels (2010) and independently as an update to Boyer et
al. (2005) for 1955–2010 (Figure 3.5a, b, e, f). Both show freshening in
the North Pacific, salinification in the North Atlantic south of 50°N and
salinification in the northern Indian Ocean (trends significant at 90%
confidence). A significant freshening is observed in the circumpolar
Southern Ocean south of 50S.
Density layers that are ventilated (connected to the sea surface) in
precipitation-dominated regions have freshened, while those venti-
lated in evaporation-dominated regions have increased in salinity,
compatible with an enhancement of the mean surface freshwater flux
pattern (Helm et al., 2010). In addition, where warming has caused
surface outcrops of density layers to move (poleward) into higher
salinity surface waters, the subducted salinity in the density layers has
increased; where outcrops have moved into fresher surface waters, the
subducted salinity decreased (Durack and Wijffels, 2010). Vertical and
lateral shifts of density surfaces, due to both changes in water mass
renewal rates and wind-driven circulation, have also contributed to
the observed subsurface salinity changes (Levitus, 1989; Bindoff and
McDougall, 1994).
A change in total, globally integrated freshwater content and salini-
ty requires an addition or removal of freshwater; the only significant
source is land ice (ice sheets and glaciers). The estimate of change in
globally averaged salinity and freshwater content remains smaller than
Figure 3.4 | (a) The 1955–2005 climatological-mean sea surface salinity (World Ocean Atlas 2009 of Antonov et al., 2010) colour contoured at 0.5 PSS78 intervals (black lines).
(b) Annual mean evaporation–precipitation averaged over the period 1950–2000 (NCEP) colour contoured at 0.5 m yr
–1
intervals (black lines). (c) The 58-year (2008 minus 1950)
sea surface salinity change derived from the linear trend (PSS78), with seasonal and El Niño-Southern Oscillation (ENSO) signals removed (Durack and Wijffels, 2010) colour con-
toured at 0.116 PSS78 intervals (black lines). (d) The 30-year (2003–2007 average centred at 2005, minus the 1960–1989 average, centred at 1975) sea surface salinity difference
(PSS78) (Hosoda et al., 2009) colour contoured at 0.06 PSS78 intervals (black lines). Contour intervals in (c) and (d) are chosen so that the trends can be easily compared, given
the different time intervals in the two analyses. White areas in (c) to (d) are marginal seas where the calculations are not carried out. Regions where the change is not significant
at the 99% confidence level are stippled in grey.
(a)
(c)
(b)
(d)
mean E -P
SSS change
60°S
30°S
0°
30°N
60°N
60°E
120°E 180°
120°W
60°W
0°
mean SSS
32 33 34 35 36 37 38
60°S
30°S
0°
30°N
60°N
60°E
120°E 180°
120°W
60°W
0°
-3 -2 -1 0 1 2 3
60°S
30°S
0°
30°N
60°N
60°E
120°E 180°
120°W
60°W
0°
-0.5 -0.4 -0.3 -0.2 -0.1 0 0.1 0.2 0.3 0.4 0.5
SSS diff
60°S
30°S
0°
30°N
60°N
60°E
120°E 180°
120°W
60°W
0°
-0.25 -0.2 -0.15 -0.1 -0.05 0 0.05 0.1 0.15 0.2 0.25
(PSS78)
(PSS78) (PSS78)
(m yr
-1
)
0
0
0
0
0
0
0
0
1
1
1
1
0
0
0
0
35
34
33
34
36
35
36
34
35
34
35
36
37
36
37
mean E -P
SSS change
269
3
Observations: Ocean Chapter 3
Frequently Asked Questions
FAQ 3.2 | Is There Evidence for Changes in the Earth’s Water Cycle?
The Earth’s water cycle involves evaporation and precipitation of moisture at the Earth’s surface. Changes in the
atmosphere’s water vapour content provide strong evidence that the water cycle is already responding to a warming
climate. Further evidence comes from changes in the distribution of ocean salinity, which, due to a lack of long-term
observations of rain and evaporation over the global oceans, has become an important proxy rain gauge.
The water cycle is expected to intensify in a warmer climate, because warmer air can be moister: the atmosphere can
hold about 7% more water vapour for each degree Celsius of warming. Observations since the 1970s show increases
in surface and lower atmospheric water vapour (FAQ 3.2, Figure 1a), at a rate consistent with observed warming.
Moreover, evaporation and precipitation are projected to intensify in a warmer climate.
Recorded changes in ocean salinity in the last 50 years support that projection. Seawater contains both salt and
fresh water, and its salinity is a function of the weight of dissolved salts it contains. Because the total amount of
salt—which comes from the weathering of rocks—does not change over human time scales, seawater’s salinity can
only be altered—over days or centuries—by the addition or removal of fresh water.
The atmosphere connects the ocean’s regions of net fresh water loss to those of fresh water gain by moving
evaporated water vapour from one place to another. The distribution of salinity at the ocean surface largely reflects
the spatial pattern of evaporation minus precipitation, runoff from land, and sea ice processes. There is some
shifting of the patterns relative to each other, because of the ocean’s currents.
Subtropical waters are highly saline, because evaporation exceeds rainfall, whereas seawater at high latitudes
and in the tropics—where more rain falls than evaporates—is less so (FAQ 3.2, Figure 1b, d). The Atlantic, the
saltiest ocean basin, loses more freshwater through evaporation than it gains from precipitation, while the Pacific
is nearly neutral (i.e., precipitation gain nearly balances evaporation loss), and the Southern Ocean (region around
Antarctica) is dominated by precipitation.
Changes in surface salinity and in the upper ocean have reinforced the mean salinity pattern. The evaporation-
dominated subtropical regions have become saltier, while the precipitation-dominated subpolar and tropical regions
have become fresher. When changes over the top 500 m are considered, the evaporation-dominated Atlantic has
become saltier, while the nearly neutral Pacific and precipitation-dominated Southern Ocean have become fresher
(FAQ 3.2, Figure 1c).
Observing changes in precipitation and evaporation directly and globally is difficult, because most of the exchange
of fresh water between the atmosphere and the surface happens over the 70% of the Earth’s surface covered
by ocean. Long-term precipitation records are available only from over the land, and there are no long-term
measurements of evaporation.
Land-based observations show precipitation increases in some regions, and decreases in others, making it difficult
to construct a globally integrated picture. Land-based observations have shown more extreme rainfall events, and
more flooding associated with earlier snow melt at high northern latitudes, but there is strong regionality in the
trends. Land-based observations are so far insufficient to provide evidence of changes in drought.
Ocean salinity, on the other hand, acts as a sensitive and effective rain gauge over the ocean. It naturally reflects
and smoothes out the difference between water gained by the ocean from precipitation, and water lost by the
ocean through evaporation, both of which are very patchy and episodic. Ocean salinity is also affected by water
runoff from the continents, and by the melting and freezing of sea ice or floating glacial ice. Fresh water added by
melting ice on land will change global-averaged salinity, but changes to date are too small to observe.
Data from the past 50 years show widespread salinity changes in the upper ocean, which are indicative of systematic
changes in precipitation and runoff minus evaporation, as illustrated in FAQ 3.2, Figure 1.
FAQ 3.2 is based on observations reported in Chapters 2 and 3, and on model analyses in Chapters 9 and 12.
(continued on next page)
270
Chapter 3 Observations: Ocean
3
FAQ 3.2, Figure 1 | Changes in sea surface salinity are related to the atmospheric patterns of evaporation minus precipitation (E – P) and trends in total precipitable
water: (a) Linear trend (1988–2010) in total precipitable water (water vapor integrated from the Earth’s surface up through the entire atmosphere) (kg m
–2
per decade)
from satellite observations (Special Sensor Microwave Imager) (after Wentz et al., 2007) (blues: wetter; yellows: drier). (b) The 1979–2005 climatological mean net E
–P (cm yr
–1
) from meteorological reanalysis (National Centers for Environmental Prediction/National Center for Atmospheric Research; Kalnay et al., 1996) (reds: net
evaporation; blues: net precipitation). (c) Trend (1950–2000) in surface salinity (PSS78 per 50 years) (after Durack and Wijffels, 2010) (blues freshening; yellows-reds
saltier). (d) The climatological-mean surface salinity (PSS78) (blues: <35; yellows–reds: >35).
31
33
35
37
(d) Mean
surface salinity
(PSS78)
−0.8
−0.4
0.0
0.4
0.8
(c) Trend in
surface salinity
(1950-2000)
(PSS78 per decade)
−100
0
100
(b) Mean
evaporation
minus
precipitation
(cm yr
-1
)
−1.6
−0.8
0.0
0.8
1.6
(a) Trend in
total precipitable
water vapour
(1988-2010)
(kg m
-2
per decade)
FAQ 3.2 (continued)
271
3
Observations: Ocean Chapter 3
its uncertainty, as was true in the AR4 assessment. For instance, the
globally averaged sea surface salinity change from 1950 to 2008 is
small (+0.003 [–0.056 to 0.062]) compared to its error estimate (Durack
and Wijffels, 2010). Thus a global freshening due to land ice loss has
not yet been discerned in global surface salinity change even if it were
assumed that all added freshwater were in the ocean’s surface layer.
3.3.3 Regional Changes in Upper Ocean Salinity
Regional changes in ocean salinity are broadly consistent with the
conclusion that regions of net precipitation (precipitation greater than
evaporation) have very likely become fresher, while regions of net
evaporation have become more saline. This pattern is seen in salinity
trend maps (Figure 3.4); zonally averaged salinity trends and freshwa-
ter inventories for each ocean (Figure 3.5); and the globally averaged
contrast between regions of high and low salinity (Figure 3.21d). In
the high-latitude regions, higher runoff, increased melting of ice and
changes in freshwater transport by ocean currents have likely also con-
tributed to observed salinity changes (Bersch et al., 2007; Polyakov et
al., 2008; Jacobs and Giulivi, 2010).
3.3.3.1 Pacific and Indian Oceans
In the tropical Pacific, surface salinity has declined by 0.1 to 0.3 over
50 years in the precipitation-dominated western equatorial regions
and by up to 0.6 to 0.75 in the Intertropical Convergence Zone and
the South Pacific Convergence Zone (Cravatte et al., 2009), while sur-
face salinity has increased by up to 0.1 over the same period in the
evaporation-dominated zones in the southeastern and north-central
tropical Pacific (Figure 3.9). The fresh, low-density waters in the warm
pool of the western equatorial Pacific expanded in area as the surface
salinity front migrated eastward by 1500 to 2500 km over the period
1955–2003 (Delcroix et al., 2007; Cravatte et al., 2009). Similarly, in the
Indian Ocean, the net precipitation regions in the Bay of Bengal and
the warm pool contiguous with the tropical Pacific warm pool have
been freshening by up to 0.1 to 0.2, while the saline Arabian Sea and
south Indian Ocean have been getting saltier by up to 0.2 (Durack and
Wijffels, 2010).
In the North Pacific, the subtropical thermocline has freshened by 0.1
since the early 1990s, following surface freshening that began around
1984 (Ren and Riser, 2010); the freshening extends down through the
intermediate water that is formed in the northwest Pacific (Nakano
et al., 2007), continuing the freshening documented by Wong et al.
(1999). Warming of the surface water that subducts to supply the inter-
mediate water is one reason for this signal, as the freshwater from the
subpolar North Pacific is now entering the subtropical thermocline at
lower density.
Salinity changes, together with temperature changes (Section 3.2.2),
affect stratification; salinity has more impact than temperature in some
regions. In the western tropical Pacific, for example, the density chang-
es from 1970 to 2003 at a trend of –0.013 kg m
–3
yr
–1
, about 60% of
that due to salinity (Delcroix et al., 2007). The decreasing density trend
mainly occurs near the surface only, which should affect stratification
across the base of the mixed layer. In the Oyashio region of the west-
ern North Pacific, salinity decrease near the surface accounts for about
60% of the density decrease of –0.004 kg m
–3
yr
–1
from 1968 to 1998
(Ono et al., 2001).
3.3.3.2 Atlantic Ocean
The net evaporative North Atlantic has become saltier as a whole over
the past 50 years (Figure 3.9; Boyer et al., 2007). The largest increase in
the upper 700 m occurred in the Gulf Stream region (0.006 per decade
between 1955–1959 and 2002–2006) (Wang et al., 2010). Salinity
increase is also evident following the circulation pathway of Mediter-
ranean Outflow Water (Figure 3.9; Fusco et al., 2008). This increase
can be traced back to the western basin of the Mediterranean, where
salinity of the deep water increased during the period from 1943 to the
mid-2000s (Smith et al., 2008; Vargas-Yáñez et al., 2010).
During the time period between 1955–1959 and 2002–2006 (using
salinities averaged over the indicated 5-year ranges), the upper 700
m of the subpolar North Atlantic freshened by up to 0.002 per decade
(Wang et al., 2010), while an increase in surface salinity was found
between the average taken over 1960–1989 and the 5-year average
over 2003–2007 (Hosoda et al., 2009). Decadal and multi-decadal
variability in the subpolar gyre and Nordic Seas is vigorous and has
been related to various climate modes such as the NAO, the Atlantic
multi-decadal oscillation (AMO, Box 2.5), and even El Niño-Southern
Oscillation (ENSO; Polyakov et al., 2005; Yashayaev and Loder, 2009),
obscuring long-term trends. The 1970s to 1990s freshening of the
northern North Atlantic and Nordic Seas (Dickson et al., 2002; Curry
et al., 2003; Curry and Mauritzen, 2005) reversed to salinification (0 to
2000 m depth) starting in the late 1990s (Boyer et al., 2007; Holliday et
al., 2008), and the propagation of this signal could be followed along
the eastern boundary from south of 60°N in the Northeast Atlantic to
Fram Strait at 79°N (Holliday et al., 2008). Advection has also played
a role in moving higher salinity subtropical waters to the subpolar
gyre (Hatun et al., 2005; Bersch et al., 2007; Lozier and Stewart, 2008;
Valdimarsson et al., 2012). The variability of the cross equatorial trans-
port contribution to this budget is highly uncertain. Reversals of North
Atlantic surface salinity of similar amplitude and duration to those
observed in the last 50 years are apparent in the early 20th century
(Reverdin et al., 2002; Reverdin, 2010). The evaporation-dominated
subtropical South Atlantic has become saltier by 0.1 to 0.3 during the
period from 1950 to 2008 (Hosoda et al., 2009; Durack and Wijffels,
2010; Figure 3.4).
3.3.3.3 Arctic Ocean
Sea ice in the Arctic has declined significantly in recent decades (Sec-
tion 4.2), which might be expected to reduce the surface salinity and
increase freshwater content as freshwater locked in multi-year sea ice
is released. Generally, strong multi-decadal variability, regional vari-
ability, and the lack of historical observations have made it difficult
to assess long-term trends in ocean salinity and freshwater content
for the Arctic as a whole (Rawlins et al., 2010). The signal that is now
emerging, including salinity observations from 2005 to 2010, indicates
increased freshwater content, with medium confidence.
Over the 20th century (1920–2003) the central Arctic Ocean in the
upper 150 m became fresher in the 1950s and then more saline by
272
Chapter 3 Observations: Ocean
3
Figure 3.5 | Zonally integrated freshwater content changes (FWCC; km
3
per degree of latitude) in the upper 500 m over one-degree zonal bands and linear trends (1955–2010)
of zonally averaged salinity (PSS78; lower panels) in the upper 500 m of the (a) and (c) Atlantic, (b) and (d) Pacific, (e) and (g) Indian and (f) and (h) World Oceans. The FWCC time
period is from 1955 to 2010 (Boyer et al., 2005; blue lines) and 1950 to 2008 (Durack and Wijffels, 2010; red lines). Data are updated from Boyer et al. (2005) and calculations of
FWCC are done according to the method of Boyer et al. (2007), using 5-year averages of salinity observations and fitting a linear trend to these averages. Error estimates are 95%
confidence intervals. The contour interval of salinity trend in the lower panels is 0.01 PSS78 per decade and dashed contours are 0.005 PSS78 per decade. Red shading indicates
values equal to or greater than 0.05 PSS78 per decade and blue shading indicates values equal to or less than –0.005 PSS78 per decade.
Freshwater content (km
3
deg
-1
)
Freshwater content (km
3
deg
-1
)
−80 −60 −40 −20 0 20 40 60 80
−1000
0
1000
2000
(a) Atlantic
Latitude
−80 −60 −40 −20 0 20 40 60 80
−1000
0
1000
2000
(b) Pacific
Latitude
−1000
0
1000
2000
(e) Indian
−1000
0
1000
2000
(f) World
−80 −60 −40 −20 020406080
−80−60 −40−20 020406080
0.01
0.01
0
0.01
0.01
0.01
0.01
0.02
0.02
0.02
0.02
0.03
0.03
0.03
0.03
0.04
-0.01
(c) Atlantic
Latitude
−80 −60 −40 −20 020406080
0
100
200
300
400
500
Depth (m)
0
0.01
0.01
0.01
0.01
-0.05
-
-0.02
-0.02
-0.01
-0.01
-0.03
-0.01
Latitude
−80−60 −40−20 020406080
(d) Pacific
0
100
200
300
400
500
0
100
200
300
400
500
0.01
0.01
0.01
0.01
0.01
-0.005
0.02
0.02
0.03
-0.01
(g) Indian
Latitude
0
100
200
300
400
500
Depth (m)
.
0.01
0.01
0.04
-0.03
-0.02
-0.01
-0.01
-0.01
(h) World
Latitude
−80 −60 −40 −20 020406080
−80−60 −40−20 020406080
273
3
Observations: Ocean Chapter 3
the early 2000s, with a net small salinification over the whole record
(Polyakov et al., 2008), while at the Siberian Shelf the river discharge
increased (Shiklomanov and Lammers, 2009) and the shelf waters
became fresher (Polyakov et al., 2008).
Upper ocean freshening has also been observed regionally in the
southern Canada basin from the period 1950–1980 to the period
1990–2000s (Proshutinsky et al., 2009; Yamamoto-Kawai et al., 2009).
These are the signals reflected in the freshwater content trend from
1955 to 2010 shown in Figure 3.5a, f: salinification at the highest lati-
tudes and a band of freshening at about 70°N to 80°N. Ice production
and sustained export of freshwater from the Arctic Ocean in response
to winds are suggested as key contributors to the high- latitude salin-
ification (Polyakov et al., 2008; McPhee et al., 2009). The contrasting
changes in different regions of the Arctic have been attributed to the
effects of Ekman transport, sea ice formation (and melt) and a shift in
the pathway of Eurasian river runoff (McPhee et al., 2009; Yamamoto-
Kawai et al., 2009; Morison et al., 2012).
Between the periods 1992–1999 and 2006–2008, not only the cen-
tral Arctic Ocean freshened (Rabe et al., 2011; Giles et al., 2012), but
also freshening is now observed in all regions including those that
were becoming more saline through the early 2000s (updated from
Polyakov et al., 2008). Moreover, freshwater transport out of the Arctic
has increased in that time period (McPhee et al., 2009).
3.3.3.4 Southern Ocean
Widespread freshening (trend of –0.01 per decade, significant at 95%
confidence, from the 1980s to 2000s) of the upper 1000 m of the
Southern Ocean was inferred by taking differences between modern
data (mostly Argo) and a long-term climatology along mean stream-
lines (Böning et al., 2008). Decadal variability, although notable, does
not overwhelm this trend (Böning et al., 2008). Both a southward shift
of the Antarctic Circumpolar Current and water-mass changes contrib-
ute to the observed trends during the period 1992–2009 (Meijers et
al., 2011). The zonally averaged freshwater content for each ocean and
the world (Figure 3.5) shows this significant Southern Ocean freshen-
ing, which exceeds other regional trends and is present in each basin
(Indian, Atlantic and Pacific, Figure 3.9).
3.3.4 Evidence for Change of the Hydrological
Cycle from Salinity Changes
The similarity between the geographic distribution of significant salin-
ity and freshwater content trends (Figures 3.4, 3.5 and 3.21) and both
the mean salinity pattern and the distribution of mean evaporation –
precipitation (E – P; Figure 3.4) indicates, with medium confidence, that
the large-scale pattern of net evaporation minus precipitation over the
oceans has been enhanced. Whereas the surface salinity pattern could
be enhanced by increased stratification due to surface warming, the
large-scale changes in column-integrated freshwater content are very
unlikely to result from changes in stratification in the thin surface layer.
Furthermore, the large spatial scale of the observed changes in fresh-
water content cannot be explained by changes in ocean circulation
such as shifts of gyre boundaries. The observed changes in surface and
subsurface salinity require additional horizontal atmospheric water
transport from regions of net evaporation to regions of net precipita-
tion. A similar conclusion was reached in AR4 (Bindoff et al., 2007). The
water vapour in the troposphere has likely increased since the 1970s,
due to warming (2.4.4, 2.5.5, 2.5.6; FAQ 3.2). The inferred enhanced
pattern of net E – P can be related to water vapor increase, although
the linkage is complex (Emori and Brown, 2005; Held and Soden,
2006). From 1950 to 2000, the large-scale pattern of surface salinity
has amplified at a rate that is larger than model simulations for the his-
torical 20th century and 21st century projections. The observed rate of
surface salinity amplification is comparable to the rate expected from
a water cycle response following the Clausius–Clapeyron relationship
(Durack et al., 2012).
Studies published since AR4, based on expanded data sets and new
analysis approaches, have substantially decreased the level of uncer-
tainty in the salinity and freshwater content trends (e.g., Stott et al.,
2008; Hosoda et al., 2009; Roemmich and Gilson, 2009; Durack and
Wijffels, 2010; Helm et al., 2010), and thus increased confidence in
the inferred changes of evaporation and precipitation over the ocean.
3.3.5 Conclusions
Both positive and negative trends in ocean salinity and freshwater con-
tent have been observed throughout much of the ocean, both at the
sea surface and in the ocean interior. While similar conclusions were
reached in AR4, the recent studies summarized here, based on expand-
ed data sets and new analysis approaches, provide high confidence in
the assessment of trends in ocean salinity. It is virtually certain that
the salinity contrast between regions of high and low surface salinity
has increased since the 1950s. It is very likely that since the 1950s,
the mean regional pattern of upper ocean salinity has been enhanced:
saline surface waters in the evaporation-dominated mid-latitudes have
become more saline, while the relatively fresh surface waters in rain-
fall-dominated tropical and polar regions have become fresher. Simi-
larly, it is very likely that the interbasin contrast between saline Atlan-
tic and fresh Pacific surface waters has increased, and it is very likely
that freshwater content in the Southern Ocean has increased. There is
medium confidence that these patterns in salinity trends are caused by
increased horizontal moisture transport in the atmosphere, suggesting
changes in evaporation and precipitation over the ocean as the lower
atmosphere has warmed.
Trends in salinity have been observed in the ocean interior as well.
It is likely that the subduction of surface water mass anomalies and
the movement of density surfaces have contributed to the observed
salinity changes on depth levels. Changes in freshwater flux and the
migration of surface density outcrops caused by surface warming (e.g.,
to regions of lower or higher surface salinity) have likely both contrib-
uted to the formation of salinity anomalies on density surfaces.
3.4 Changes in Ocean Surface Fluxes
3.4.1 Introduction
Exchanges of heat, water and momentum (wind stress) at the sea sur-
face are important factors for driving the ocean circulation. Changes
274
Chapter 3 Observations: Ocean
3
in the air–sea fluxes may result from variations in the driving surface
meteorological state variables (air temperature and humidity, SST,
wind speed, cloud cover, precipitation) and can impact both water-
mass formation rates and ocean circulation. Air–sea fluxes also influ-
ence temperature and humidity in the atmosphere and, therefore, the
hydrological cycle and atmospheric circulation. AR4 concluded that,
at the global scale, the accuracy of the observations is insufficient to
permit a direct assessment of changes in heat flux (AR4 Section 5.2.4).
As described in Section 3.4.2, although substantial progress has been
made since AR4, that conclusion still holds for this assessment.
The net air–sea heat flux is the sum of two turbulent (latent and sensi-
ble) and two radiative (shortwave and longwave) components. Ocean
heat gain from the atmosphere is defined to be positive according to
the sign convention employed here. The latent and sensible heat fluxes
are computed from the state variables using bulk parameterizations;
they depend primarily on the products of wind speed and the verti-
cal near-sea-surface gradients of humidity and temperature respec-
tively. The air–sea freshwater flux is the difference of precipitation (P)
and evaporation (E). It is linked to heat flux through the relationship
between evaporation and latent heat flux. Thus, when considering
potential trends in the global hydrological cycle, consistency between
observed heat budget and evaporation changes is required in areas
where evaporation is the dominant term in hydrological cycle changes.
Ocean surface shortwave and longwave radiative fluxes can be inferred
from satellite measurements using radiative transfer models, or com-
puted using empirical formulae, involving astronomical parameters,
atmospheric humidity, cloud cover and SST. The wind stress is given by
the product of the wind speed squared, and the drag coefficient. For
detailed discussion of all terms see, for example, Gulev et al. (2010).
Atmospheric reanalyses, discussed in Box 2.3, are referred to fre-
quently in the following sections and for clarity the products cited are
summarised here: ECMWF 40-year Reanalysis (referred to as ERA40
hereafter, Uppala et al., 2005), ECMWF Interim Reanalysis (ERAI, Dee
et al., 2011), NCEP/NCAR Reanalysis 1 (NCEP1, Kalnay et al., 1996),
NCEP/DOE Reanalysis 2 (NCEP2, Kanamitsu et al., 2002), NCEP Climate
Forecast System Reanalysis (CFSR, Saha et al., 2010), NASA Modern
Era Reanalysis for Research and Applications (MERRA, Rienecker et al.,
2011) and NOAA-CIRES 20th Century Reanalysis, version 2 (20CRv2,
Compo et al., 2011).
3.4.2 Air–Sea Heat Fluxes
3.4.2.1 Turbulent Heat Fluxes and Evaporation
The latent and sensible heat fluxes have a strong regional dependence,
with typical values varying in the annual mean from close to zero to
–220 W m
–2
and –70 W m
–2
respectively over strong heat loss sites (Yu
and Weller, 2007). Estimates of these terms have many potential sourc-
es of error (e.g., sampling issues, instrument biases, changing data
sources, uncertainty in the flux computation algorithms). These sources
may be spatially and temporally dependent, and are difficult to quan-
tify (Gulev et al., 2007); consequently flux error estimates have a high
degree of uncertainty. Spurious temporal trends may arise as a result
of variations in measurement method for the driving meteorological
state variables, in particular wind speed (Tokinaga and Xie, 2011). The
overall uncertainty of the annually averaged global ocean mean for
each term is expected to be in the range 10 to 20%. In the case of the
latent heat flux term, this corresponds to an uncertainty of up to 20 W
m
–2
. In comparison, changes in global mean values of individual heat
flux components expected as a result of anthropogenic climate change
since 1900 are at the level of <2 W m
–2
(Pierce et al., 2006).
Many new turbulent heat flux data sets have become available since
AR4 including products based on atmospheric reanalyses, satellite and
in situ observations, and hybrid or synthesized data sets that combine
information from these three different sources. It is not possible to
identify a single best product as each has its own strengths and weak-
nesses (Gulev et al., 2010); several data sets are summarised here to
illustrate the key issues. The Hamburg Ocean-Atmosphere Parameters
and Fluxes from Satellite (HOAPS) data product provides global tur-
bulent heat fluxes (and precipitation) developed from observations
at microwave and infrared wavelengths (Andersson et al., 2011). In
common with other satellite data sets it provides globally complete
fields, however, it spans a relatively short period (1987 onwards) and
is thus of limited utility for identifying long-term changes. A significant
advance in flux data set development methodology is the 1 × 1 degree
grid Objectively Analysed Air–Sea heat flux (OAFlux) data set that
covers 1958 onwards and for the first time synthesizes state variables
(SST, air temperature and humidity, wind speed) from reanalyses and
satellite observations, prior to flux calculation (Yu and Weller, 2007).
OAFlux has the potential to minimize severe spatial sampling errors
that limit the usefulness of data sets based on ship observations alone
and provides a new resource for temporal variability studies. However,
the data sources for OAFlux changed in the 1980s, with the advent of
satellite data, and the consequences of this change need to be assessed.
In an alternative approach,Large and Yeager (2009) modified NCEP1
reanalysis state variables prior to flux calculation using various adjust-
ment techniques, to produce the hybrid Coordinated Ocean-ice Refer-
ence Experiments (CORE) turbulent fluxes for 1948–2007(Griffies et
al., 2009). However, as the adjustments employed to produce the CORE
fluxes were based on limited periods (e.g., 2000–2004 for wind speed)
it is not clear to what extent CORE can be reliably used for studies of
interdecadal variability over the 60-year period that it spans.
Analysis of OAFlux suggests that global mean evaporation may vary
at inter-decadal time scales, with the variability being relatively small
compared to the mean (Yu, 2007; Li et al., 2011; Figure 3.6a). Changing
data sources, particularly as satellite observations became available in
the 1980s, may contribute to this variability (Schanze et al., 2010) and
it is not yet possible to identify how much of the variability is due to
changes in the observing system. The latent heat flux variations (Figure
3.6b) closely follow those in evaporation (with allowance for the sign
definition which results in negative values of latent heat flux corre-
sponding to positive values of evaporation) but do not scale exactly
as there is an additional minor dependence on SST through the latent
heat of evaporation. The large uncertainty ranges that are evident in
each of the time series highlight the difficulty in establishing whether
there is a trend in global ocean mean evaporation or latent heat flux.
The uncertainty range for latent heat flux is much larger than the 0.5 W
m
–2
level of net heat flux change expected from the ocean heat content
increase (Box 3.1). Thus, it is not yet possible to use such data sets to
establish global ocean multi-decadal trends in evaporation or latent
275
3
Observations: Ocean Chapter 3
heat flux at this level. The globally averaged sensible heat flux is small-
er in magnitude than the latent heat flux and has a smaller absolute
range of uncertainty (Figure 3.6b).
3.4.2.2 Surface Fluxes of Shortwave and Longwave Radiation
The surface shortwave flux has a strong latitudinal dependence with
typical annual mean values of 250 W m
–2
in the tropics. The annual mean
surface net longwave flux ranges from –30 to –70 W m
–2
. Estimates of
these terms are available from in situ climatologies, from atmospher-
ic reanalyses, and, since the 1980s, from satellite observations. These
data sets have many potential sources of error that include: uncer-
tainty in the satellite retrieval algorithms and in situ formulae, cloud
representation in reanalyses, sampling issues and changing satellite
sensors (Gulev et al., 2010). As for the turbulent fluxes, the uncertainty
of the annually averaged global ocean mean shortwave or longwave
flux is difficult to determine and in the range 10–20%.
High accuracy in situ radiometer measurements are available at land
sites since the 1960s (see Wild, 2009 Figure 1), allowing analysis of
decadal variations in the surface shortwave flux. However, this is not
the case over the oceans, where there are very few in situ measure-
ments (the exception being moored buoy observations in the tropical
band 15°S to 15°N since the 1990s, Pinker et al., 2009). Consequently,
for global ocean shortwave analyses it is necessary to rely on satellite
observations, which are less accurate (compared to in situ determi-
nation of radiative fluxes), restrict the period that can be considered
to the mid-1980s onwards, but do provide homogeneous sampling.
Detailed discussion of variations in global (land and ocean) averaged
surface solar radiation is given in Section 2.3.3; confidence in variabili-
ty of radiation averaged over the global ocean is low owing to the lack
of direct observations.
3.4.2.3 Net Heat Flux and Ocean Heat Storage Constraints
The most reliable source of information for changes in the global mean
net air–sea heat flux comes from the constraints provided by analyses
of changes in ocean heat storage. The estimate of increase in global
ocean heat content for 1971–2010 quantified in Box 3.1 corresponds
to an increase in mean net heat flux from the atmosphere to the ocean
of 0.55 W m
–2
. In contrast, closure of the global ocean mean net sur-
face heat flux budget to within 20 W m
–2
from observation based sur-
face flux data sets has still not been reliably achieved (e.g., Trenberth et
al., 2009). The increase in mean net air–sea heat flux is thus small com-
pared to the uncertainties of the global mean. Large and Yeager (2012)
examined global ocean average net heat flux variability using the
COREdata set over 1984–2006 and concluded thatnatural variability,
rather than long-term climate change, dominates heat flux changes
over this relatively short, recent period. Since AR4, some studies have
shown consistency in regional net heat flux variability at sub-basin
scale since the 1980s, notably in the Tropical Indian Ocean (Yu et al.,
2007) and North Pacific (Kawai et al., 2008). However, detection of a
change in air–sea fluxes responsible for the long-term ocean warming
remains beyond the ability of currently available surface flux data sets.
3.4.3 Ocean Precipitation and Freshwater Flux
Assessment of changes in ocean precipitation at multi-decadal time
scales is very difficult owing to the lack of reliable observation based
data sets prior to the satellite era. The few studies available rely
on reconstruction techniques. Remote sensing based precipitation
observations from the Global Precipitation Climatology Project (GPCP)
for 1979–2003 have been used by Smith et al. (2009, 2012) to recon-
struct precipitation for 1900–2008 (over 75°S to 75°N) by employ-
ing statistical techniques that make use of the correlation between
precipitation and both SST and sea level pressure (SLP). Each of the
reconstructions shows both centennial and decadal variability in global
ocean mean precipitation (Figure 3.7). The trend from 1900 to 2008 is
1.5 mm per month
per century according to Smith et al. (2012). For the
period of overlap, the reconstructed global ocean mean precipitation
Figure 3.6 | Time series of annual mean global ocean average evaporation (red line,
a), sensible heat flux (green line, b) and latent heat flux (blue line, b) from 1958 to 2012
determined by Yu from a revised and updated version of the original OAFlux data set Yu
and Weller (2007). Shaded bands show uncertainty estimates and the black horizontal
bars in (b) show the time periods for which reanalysis output and satellite observations
were employed in the OAFlux analysis; they apply to both panels.
1960
1970
1980
1990
2000
2010
-50
-30
-10
-100
-80
-60
-40
-20
0
Year
Latent and sensible heat (W m
-2
)
(b)
1960
1970
1980
1990
2000
2010
90
95
100
105
110
115
120
125
130
Year
Evaporation (cm yr
-1
)
(a)
Reanalysis
Satellite
-70
-90
-110
Sensible
Latent
276
Chapter 3 Observations: Ocean
3
time series show consistent variability with GPCP as is to be expected
(Figure 3.7). Focusing on the Tropical Ocean (25°S to 25°N) for the
recent period 1979–2005, Gu et al. (2007) have identified a precipi-
tation trend of 0.06 mm day
–1
per decade using GPCP. Concerns have
been expressed in the cited studies over the need for further work both
to determine the most reliable approach to precipitation reconstruc-
tion and to evaluate the remotely sensed precipitation data sets. Given
these concerns, confidence in ocean precipitation trend results is low.
Evaporation and precipitation fields from atmospheric reanalyses
can be tested for internal consistency of different components of the
hydrological cycle. Specifically, the climatological mean value for E – P
averaged over the global ocean should equal both the correspond-
ing mean for P – E averaged over land and the moisture transport
from ocean to land. Trenberth et al. (2011) find in an assessment of
eight atmospheric reanalyses that this is not the case for each product
considered, and they also report spurious trends due to variations in
the observing system with time. Schanze et al. (2010) examine interan-
nual variability within the OAFlux evaporation and GPCP precipitation
data sets, and find that use of satellite data prior to 1987 is limited by
discontinuities attributable to variations in data type. Thus, it is not yet
possible to use such data sets to establish whether there are significant
multi-decadal trends in mean E – P. However, regional trends in surface
salinity since the 1950s do suggest trends in E – P over the same time
(see Section 3.3.4).
3.4.4 Wind Stress
Wind stress fields are available from reanalyses, satellite-based data
sets, and in situ observations. Basin scale wind stress trends at decadal
to centennial time scales have been reported for the Southern Ocean,
the North Atlantic and the Tropical Pacific as detailed below. Howev-
er, these results are based largely on atmospheric reanalyses, in some
cases a single product, and consequently the confidence level is low to
medium depending on region and time scale considered.
In the Southern Ocean, the majority of reanalyses in the most compre-
hensive study available show an increase in the annual mean zonal
wind stress (Swart and Fyfe, 2012; Figure 3.8). They find an increase
in annual mean wind stress strength in four (NCEP1, NCEP2, ERAI and
20CRv2) of the six reanalyses considered (Figure 3.8). The mean of all
reanalyses available at a given time (Figure 3.8, black line) also shows
an upward trend from about 0.15 N m
–2
in the early 1950s to 0.20 N
m
–2
in the early 2010s. An earlier study, covering 1979–2009, found a
wind stress increase in two of four reanalyses considered (Xue et al.,
2010). A positive trend of zonal wind stress from 1980 to 2000 was
also reported by Yang et al. (2007) using a single reanalysis (ERA40)
and found to be consistent with increases in wind speed observations
made on Macquarie Island (54.5°S, 158.9°E) and by the SSM/I satel-
lite (data from 1987 onwards). The wind stress strengthening is found
by Yang et al. (2007) to have a seasonal dependence, with strongest
trends in January, and has been linked by them to changes in the
Southern Annular Mode (SAM, Box 2.5), which has continued to show
an upward trend since AR4 (Section 2.7.8). Taken as a whole, these
studies provide medium confidence that Southern Ocean wind stress
has strengthened since the early 1980s. A strengthening of the related
wind speed field in the Southern Ocean, consistent with the increasing
trend in the SAM, has also been noted in Section 2.7.2 from satel-
lite-based analyses and atmospheric reanalyses.
In the Tropical Pacific, a reanalysis based study found a strengthening
of the trade wind associated wind stress for 1990–2009, but for the
earlier period 1959–1989 there is no clear trend (Merrifield, 2011).
Strengthening of the related Tropical Pacific Ocean wind speed field
in recent decades is evident in reanalysis and satellite based data sets.
Taken together with evidence for rates of sea level rise in the western
Pacific larger than the global mean (Section 3.7.3) these studies pro-
vide medium confidence that Tropical Pacific wind stress has increased
since 1990. This increase may be related to the Pacific Decadal Oscilla-
tion (Merrifield et al., 2012). At centennial time scales, attempts have
been made to reconstruct the wind stress field in the Tropical Pacific by
making use of the relationship between wind stress and SLP in combi-
nation with historic SLP data. Vecchi et al. (2006), using this approach,
found a reduction of 7% in zonal mean wind stress across the Equa-
torial Pacific from the 1860s to the 1990s and related it to a possible
weakening of the tropical Walker circulation. Observations discussed in
Section 2.7.5 indicate that this weakening has largely been offset by a
stronger Walker circulation since the 1990s.
Changes in winter season wind stress curl over the North Atlantic
from 1950 to early 2000s from NCEP1 and ERA40 have leading modes
that are highly correlated with the NAO and East Atlantic circulation
patterns; each of these patterns demonstrates a trend towards more
Figure 3.7 | Long-term reconstruction of ocean precipitation anomaly averaged over
75°S to 75°N from Smith et al. (2012): Annual values, thin blue line; low-pass filtered
(15-year running mean) values, bold blue line with uncertainty estimates (shading).
Smith et al. (2009) low-pass filtered values, dotted grey line. Also shown is the cor-
responding GPCPv2.2 derived ocean precipitation anomaly time series averaged over
the same latitudinal range (annual values, thin magenta line; low-pass filtered values,
bold magenta line); note Smith et al. (2012) employed an earlier version of the GPCP
data set leading to minor differences relative to the published time series in their paper.
Precipitation anomalies were taken relative to the 1979–2008 period.
1900191019201930 1940 1950 1960 1970 1980 1990 2000 2010
Year
-4
-3
-2
-1
0
1
2
3
Precipitation anomaly (mm per month)
Smith et al. (2012)
GPCP
Smith et al. (2009)
277
3
Observations: Ocean Chapter 3
positive index values superimposed on pronounced decadal variability
over the period from the early 1960s to the late 1990s (Sugimoto and
Hanawa, 2010). Wu et al. (2012) find a poleward shift over the past
century of the zero wind stress curl line by 2.5° [1.5° to 3.5°] in the
North Atlantic and 3.0° [1.6° to 4.4°] in the North Pacific from 20CRv2.
Confidence in these results is low as they are based on a single prod-
uct, 20CRv2 (the only century time scale reanalysis), which may be
affected by temporal inhomogeneity in the number of observations
assimilated (Krueger et al., 2013).
3.4.5 Changes in Surface Waves
Surface wind waves are generated by wind forcing and are partitioned
into two components, namely wind–sea (wind-forced waves propagat-
ing slower than surface wind) and swell (resulting from the wind–sea
development and propagating typically faster than surface wind). Sig-
nificant wave height (SWH) represents the measure of the wind wave
field consisting of wind–sea and swell and is approximately equal to
the highest one-third of wave heights. Local wind changes influence
wind–sea properties, while changes in remote storms affect swell.
Thus, patterns of wind wave and surface wind variability may differ
because wind waves integrate wind properties over a larger domain.
As wind waves integrate characteristics of atmospheric dynamics over
a range of scales they potentially serve as an indicator of climate var-
iability and change. Global and regional time series of wind waves
characteristics are available from buoy data, Voluntary Observing Ship
(VOS) reports, satellite measurements and model wave hindcasts. No
source is superior, as all have their strengths and weaknesses (Sterl
and Caires, 2005; Gulev and Grigorieva, 2006; Wentz and Ricciardulli,
2011).
3.4.5.1 Changes in Surface Waves from Voluntary Observing
Ship and Wave Model Hindcasts Forced by Reanalyses
AR4 reported statistically significant positive SWH trends during 1900–
2002 in the North Pacific (up to 8 to 10 cm per decade) and strong-
er trends (up to 14 cm per decade) from 1950 to 2002 for most of
the mid-latitudinal North Atlantic and North Pacific, with insignificant
trends, or small negative trends, in most other regions (Trenberth et
al., 2007). Studies since AR4 have provided further evidence for SWH
trends with more detailed quantification and regionalization.
Model hindcasts based on 20CRv2 (spanning 1871–2010) and ERA40
(spanning 1958–2001) show increases in annual and winter mean
SWH in the north-east Atlantic, although the trend magnitudes depend
on the reanalysis products used (Sterl and Caires, 2005; Wang et al.,
2009, 2012; Semedo et al., 2011). Analysis of VOS observations for
1958–2002 reveals increases in winter mean SWH over much of the
North Atlantic, north of 45°N, and the central to eastern mid-latitude
North Pacific with typical trends of up to 20 cm per decade (Gulev and
Grigorieva, 2006).
3.4.5.2 Changes in Surface Waves from Buoy Data
Positive regional trends in extreme wave heights have been reported
at several buoy locations since the late 1970s, with some evidence for
seasonal dependence, including at sites on the east and west coasts of
the USA (Komar and Allan, 2008; Ruggiero et al., 2010) and the north-
east Pacific coast (Menéndez et al., 2008). However, Gemmrich et al.
(2011) found for the Pacific buoys that some trends may be artefacts
due to step-type historical changes in the instrument types, observa-
tional practices and post-processing procedures. Analysis of data from
a single buoy deployed west of Tasmania showed no significant trend
in the frequency of extreme waves contrary to a significant positive
trend seen in the ERA40 reanalysis (Hemer, 2010).
3.4.5.3 Changes in Surface Waves from Satellite Data
Satellite altimeter observations provide a further data source for wave
height variability since the mid-1980s. Altimetry is of particular value
in the southern hemisphere, and in some poorly sampled regions of
the northern hemisphere, where analysis of SWH trends remains a
challenge due to limited in situ data and temporal inhomogeneity in
the data used for reanalysis products. In the Southern Ocean, altime-
ter-derived SWH and model output both show regions with increas-
ing wave height although these regions cover narrower areas in the
altimeter analysis than in the models and have smaller trends (Hemer
et al., 2010). Young et al. (2011a) compiled global maps of mean and
extreme (90th and 99th percentile) surface wind speed and SWH
trends for 1985–2008 using altimeter measurements. As the length
of the data set is short, it is not possible to determine whether their
results reflect long-term SWH and wind speed trends, or are part of
a multi-decadal oscillation. For mean SWH, their analysis shows pos-
itive linear trends of up to 10 to 15 cm per decade in some parts of
the Southern Ocean (with the strongest changes between 80°E and
160°W) that may reflect the increase in strength of the wind stress
since the early 1980s (see Section 3.4.4). Young et al. (2011a) note,
however, that globally the level of statistical significance is generally
low in the mean and 90th percentile SWH trends but increases for the
99th percentile. Small negative mean SWH trends are found in many
NH ocean regions and these are of opposite sign to, and thus incon-
sistent with, trends in wind speed — the latter being primarily positive.
Nevertheless, for the 99th SWH percentile, strong positive trends up to
50 to 60 cm per decade were identified in the Southern Ocean, North
Figure 3.8 | Time series of annual average maximum zonal-mean zonal wind stress
(N m
–2
) over the Southern Ocean for various atmospheric reanalyses: CFSR (orange),
NCEP1 (cyan), NCEP2 (red), ERAI (dark blue), MERRA (green), 20CR (grey), and mean
of all reanalyses at a given time (thick black), see Box 2.3 for details of reanalyses.
Updated version of Figure 1a in Swart and Fyfe (2012), with CFSR, MERRA and the
mean of all reanalyses added.
1950
1960
1970
1980
1990
2000
2010
0.125
0.15
0.175
0.2
0.225
0.25
Year
Wind stress (N m )
-2
278
Chapter 3 Observations: Ocean
3
Atlantic and North Pacific and these are consistent in sign with the
extreme wind speed trends. Subsequent analysis has shown that the
Young et al. (2011a) wind speed trends tend to be biased high when
compared with microwave radiometer data (Wentz and Ricciardulli,
2011; Young et al., 2011b).
3.4.6 Conclusions
Uncertainties in air–sea heat flux data sets are too large to allow detec-
tion of the change in global mean net air–sea heat flux, on the order
of 0.5 W m
–2
since 1971, required for consistency with the observed
ocean heat content increase. The accuracy of reanalysis and satellite
observation based freshwater flux products is limited by changing data
sources. Consequently, the products cannot yet be reliably used to
directly identify trends in the regional or global distribution of evapora-
tion or precipitation over the oceans on the time scale of the observed
salinity changes since 1950.
Basin scale wind stress trends at decadal to centennial time scales
have been observed in the North Atlantic, Tropical Pacific, and South-
ern Oceans with low to medium confidence. These results are based
largely on atmospheric reanalyses, in some cases a single product, and
the confidence level is dependent on region and time scale consid-
ered. The evidence is strongest for the Southern Ocean for which there
is medium confidence that zonal mean wind stress has increased in
strength since the early 1980s.
There is medium confidence based on ship observations and reanalysis
forced wave model hindcasts that mean significant wave height has
increased since the 1950s over much of the North Atlantic north of
45°N, with typical winter season trends of up to 20 cm per decade.
3.5 Changes in Water-Mass Properties
3.5.1 Introduction
To a large degree, water properties are set at the sea surface through
interaction between the ocean and the overlying atmosphere (and ice,
in polar regions). The water characteristics resulting from these interac-
tions (e.g., temperature, salinity and concentrations of dissolved gases
and nutrients) are transferred to various depths in the world ocean,
depending on the density of the water. Warm, light water masses
supply (or “ventilate”) the upper ocean at low to mid-latitudes, while
the colder, denser water masses formed at higher latitudes supply the
intermediate and deep layers of the ocean (see schematic in FAQ 3.1,
Figure 1). The formation and subduction of water masses are important
for the ocean’s capacity to store heat, freshwater, carbon, oxygen and
other properties relevant to climate. In this section, the evidence for
change in some of the major water masses of the world ocean is
assessed.
The zonal-mean distributions of salinity, density, and temperature in
each ocean basin (black contours in Figure 3.9) reflect the formation
of water masses at the sea surface and their subsequent spreading
into the ocean interior. For example, warm, salty waters formed in the
regions of net evaporation between 10° and 30° latitude (Figure 3.4b)
supply the subtropical salinity maximum waters found in the upper
few hundred meters in each basin (Figure 3.9). Relatively fresh water
masses produced at higher latitude, where precipitation exceeds evap-
oration, sink and spread equatorward to form salinity minimum layers
at intermediate depths. Outflow of saline water from the Mediterrane-
an Sea and Red Sea, where evaporation is very strong, accounts for the
relatively high salinity observed in the upper 1000 m in the subtropical
North Atlantic and North Indian basins, respectively.
Many of the observed changes in zonally averaged salinity, density and
temperature are aligned with the spreading paths of the major water
masses (Figure 3.9, trends from 1950 to 2000 shown in colours and
white contours), illustrating how the formation and spreading of water
masses transfer anomalies in surface climate to the ocean interior. The
strongest anomalies in a water mass are found near its source region.
For instance, bottom and deep water anomalies are strongest in the
Southern Ocean and the northern North Atlantic, with lessening ampli-
tudes along the spreading paths of these water masses. In each basin,
the subtropical salinity maximum waters have become more saline,
while the low-salinity intermediate waters have become fresher (Figure
3.9 a, d, g, j; see also Section 3.3). Strongest warming is observed in
the upper 100 m, which has warmed almost everywhere, with reduced
warming (Atlantic) or regions of cooling (Indian and Pacific) observed
between 100 and 500 m depth.
Warming is observed throughout the upper 2000 m south of 40°S in
each basin. Shifts in the location of ocean circulation features can also
contribute to the observed trends in temperature and salinity, as dis-
cussed in Section 3.2. Density decreased throughout most of the upper
2000 m of the global ocean (middle column of Figure 3.9). The decrease
in near-surface density (hence increase in stratification) is largest in the
Pacific, where warming and freshening both act to reduce density, and
smallest in the Atlantic where the salinity and temperature trends have
opposite effects on density.
The remainder of this section focuses on evidence of change in globally
relevant intermediate, deep and bottom water masses.
3.5.2 Intermediate Waters
3.5.2.1 North Pacific Intermediate Water
The North Pacific Intermediate Water (NPIW) has freshened over
the last two decades (Wong et al., 1999; Nakano et al., 2007; Figure
3.9g) and has warmed since the 1950s, as reported in AR4, Chapter 5.
NPIW in the northwestern North Pacific warmed by 0.5°C from 1955
to 2004 and is now entering the subtropics at lower density; oxygen
concentrations in the NPIW have declined, indicating weaker ventila-
tion (Nakanowatari et al., 2007; Kouketsu et al., 2010). The strongest
trends are in the Sea of Okhotsk, where NPIW is formed, and have
been tentatively linked to increased air temperature and decreased
sea-ice extent in winter (Nakanowatari et al., 2007; Figure 3.9i).
3.5.2.2 Antarctic Intermediate Water
In AR4, Chapter 5, Antarctic Intermediate Water (AAIW) was reported
to have warmed and freshened since the 1960s (Figure 3.9). In most
279
3
Observations: Ocean Chapter 3
recent studies, usually—but not always—a dipole pattern was found:
on isopycnals denser than the AAIW salinity minimum, a warming and
salinification was observed and on isopycnals lighter than the AAIW
salinity minimum, a cooling and freshening trend (Böning et al., 2008;
Durack and Wijffels, 2010; Helm et al., 2010; McCarthy et al., 2011).
The salinity minimum core of the AAIW also underwent changes con-
sistent with these patterns on isopycnals: In 1970–2009, south of 30°S,
the AAIW salinity minimum core showed a strong, large-scale shoal-
ing (30 to 50 dbar per decade) and warming (0.05°C to 0.15°C per
decade), leading to lighter densities (up to 0.03 kg m
–3
per decade),
while the salinity trends varied regionally. A long-term freshening of
the AAIW core is found in the southwest Atlantic, southeast Pacific,
and south-central Indian oceans, with salinification south of Africa
and Australia. All trends were strongest close to the AAIW formation
latitude just north of the Antarctic Circumpolar Current (Schmidtko and
Johnson, 2012).
Both an increase in precipitation—evaporation and poleward migra-
tion of density surfaces caused by warming have likely contributed
to the observed trends (Section 3.3; Böning et al., 2008; Durack and
Wijffels, 2010; Helm et al., 2010; McCarthy et al., 2011). Changes in
AAIW properties in particular locations have also been linked to other
processes, including exchange between the Indian and Atlantic basins
(McCarthy et al., 2011) and changes in surface forcing related to modes
of climate variability like ENSO and the SAM (Garabato et al., 2009).
Whether these changes in properties also affected the formation rates
of AAIW cannot be assessed from the available observations.
3.5.3 Deep and Bottom Waters
Deep and bottom layers of the ocean are supplied by roughly equal
volumes of dense water sinking in the northern North Atlantic (Lower
North Atlantic Deep Water, LNADW) and around Antarctica (Antarctic
Bottom Water, AABW) (FAQ 3.1, Figure 1).
3.5.3.1 Upper North Atlantic Deep Water
Upper North Atlantic Deep Water (UNADW) is formed by deep convec-
tion in the Labrador Sea between Canada and Greenland, so is also
known as Labrador Sea Water (LSW). It is the shallowest component of
the NADW, located above the overflow water masses that supply the
Lower North Atlantic Deep Water (LNADW). AR4 Chapter 5 assessed
the variability in water mass properties of LSW from the 1950s. Recent
studies have confirmed the large interannual-to-multi-decadeal varia-
bility of LSW properties and provided new information on variability
in formation rates and the impact on heat and carbon (Section 3.8.1)
uptake by the deep ocean.
During the 1970s and 1980s and especially the 1990s the UNADW has
been cold and fresh. In Figure 3.9A it is the strong freshening signal
from the 1960s to the 1990s that dominates the trend. This freshen-
ing trend reversed in the late 1990s (Boyer et al., 2007; Holliday et
al., 2008; see Section 3.3.3.2). Estimates of the LSW formation rate
3
decreased from about 7.6 to 8.9 Sv in 1997–1999 (Kieke et al., 2006) to
roughly 0.5 Sv in 2003–2005 (Rhein et al., 2011), and since 1997, only
less dense LSW was formed compared to the high NAO years before.
There is, however, evidence that formation of denser LSW occurred in
2008 (Våge et al., 2009; Yashayaev and Loder, 2009), but not in the
following years (Yashayaev and Loder, 2009; Rhein et al., 2011).
The strong variability in the formation of UNADW affected significantly
the heat transfer into the deep North Atlantic (Mauritzen et al., 2012).
Substantial heat entered the deep North Atlantic during the low NAO
years of the 1960s, when salinity was large enough to compensate for
the high temperatures, and dense LSW was still formed and exported
to the subtropics.
3.5.3.2 Lower North Atlantic Deep Water
Dense waters overflowing the sills between Greenland and Scotland
supply the Lower North Atlantic Deep Water (LNADW). Both overflows
freshened from the mid-1960s to the mid-1990s (Dickson et al., 2008).
The salinity of the Faroe Bank overflow increased by 0.015 to 0.02
from 1997 to 2004, implying a density increase on the order of 0.01
kg m
–3
(Hansen and Osterhus, 2007). The other main overflow, through
Denmark Strait, shows large interannual variability in temperature
and salinity, but no trends for the time period 1996–2011 (Jochum-
sen et al., 2012). Observations of the transport of the dense overflows
are dominated by short-term variability and there is no evidence of
a trend in the short time series available (see Section 3.6). As both
overflow components descend into the North Atlantic, they entrain
substantial amounts of ambient subpolar waters to create LNADW. As
a whole, the LNADW in the North Atlantic cooled from the 1950s to
2005 (Mauritzen et al., 2012), a signal thus stemming primarily from
the entrained waters, possibly an adjustment from an unusually warm
period observed in the 1920s and 1930s (Drinkwater, 2006).
3.5.3.3 Antarctic Bottom Water
The Antarctic Bottom Water (AABW) has warmed since the 1980s
or 1990s, most noticeably near Antarctica (Aoki et al., 2005; Rintoul,
2007; Johnson et al., 2008a; Purkey and Johnson, 2010; Kouketsu et
al., 2011), but with warming detectable into the North Pacific and
North Atlantic Oceans (Johnson et al., 2008b; Kawano et al., 2010). The
warming of AABW between the 1990s and 2000s contributed to global
ocean heat uptake (Section 3.2). The global volume of the AABW layer
decreased by 8.2 [5.6 to 10.8] Sv during the last two decades (John-
son et al., 2008b; Mauritzen et al., 2012; Purkey and Johnson, 2012),
making it more likely than not that at least the export rate of AABW
from the Southern Ocean declined during this period.
The sources of AABW in the Indian and Pacific sectors of the Southern
Ocean have freshened in recent decades. The strongest signal (0.03
per decade, between 1970 and 2008) is observed in the Ross Sea and
has been linked to inflow of glacial melt water from the Amundsen
and Bellingshausen Seas (Shepherd et al., 2004; Rignot et al., 2008;
Jacobs and Giulivi, 2010). Freshening has been observed in AABW
since the 1970s in the Indian sector (Rintoul, 2007) and between the
3
The formation rate of a water mass is the volume of water per year that is transformed into the density range of this water mass by surface processes (for instance cooling),
eventually modified through ocean interior processes (for instance mixing). Formation rates are reported in Sverdrups (Sv). 1 Sv equals 10
6
m
3
s
–1
.
280
Chapter 3 Observations: Ocean
3
1990s and 2000s in the Pacific sector (Swift and Orsi, 2012; Purkey
and Johnson, 2013).
In the Weddell Sea (the primary source of AABW in the Atlantic), a
contraction of the bottom water mass was observed between 1984
and 2008 at the Prime Meridian, accompanied by warming of about
0.015°C, and by salinity variability on a multi-annual time scale. Tran-
sient tracer observations between 1984 and 2011 confirmed that the
AABW there has become less well ventilated over that time period. The
changes in the AABW, however, seem to be caused by the much strong-
er trends observed in the Warm Deep Water, as WDW is entrained into
the AABW while sinking to the bottom, and not by changes in the
AABW formation rate (Huhn et al., 2008; Huhn et al., 2013).
3.5.4 Conclusions
AR4 Chapter 5 concluded that observed changes in upper ocean water
masses reflect the combination of long-term trends and interannual to
decadal variability related to climate modes like ENSO, NAO and SAM.
The time series are still generally too short and incomplete to distinguish
decadal variability from long-term trends, but understanding of the
nature and causes of variability has improved in this assessment. The
observed patterns of change in subsurface temperature and salinity
(Sections 3.2 and 3.3) are consistent with understanding of how and
where water masses form, enhancing the level of confidence in the
assessment of the observed changes.
3
5
35
Pressure (dbar)
0
100
200
300
400
34
.5
Latitude
Salinity
J
GLO
70S 50S 30S 10S 10N 30N 50N 70N
500
1000
1500
2000
−
0
.2 −
0
.1
5
−
0
.1 −
0
.
05 0 0
.
05 0
.1
0
.1
5 0
.2
23
24
25
26
27
2
7
28
28
Latitude
Density
K
50S 30S 10S 10N 30N 50N 70N
−
0
.
3
−
0
.2 −
0
.1
0 0
.1
0
.2
0
.
3
0
5
5
5
1
0
10
15
20
25
0
5
Latitude
Temperature
L
50S 30S 10S 10N 30N 50N 70N
−1 −
0
.7
5
−
0
.
5
−
0
.2
5 0 0
.2
5 0
.
5 0
.7
5
1
34
35
Pressure (dbar)
0
100
200
300
400
34.5
34.5
G
PAC
500
1000
1500
2000
23
24
25
2
6
2
7
H
5
5
10
1
5
20
25
5
I
35
35
35
35.5
36
36
Pressure (dbar)
0
100
200
300
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34.
5
35
35
A
ATL
500
1000
1500
2000
2
4
25
26
27
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28
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2
7
B
0
5
5
10
10
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5
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0
0
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10
C
3
5
35
36
Pressure (dbar)
0
100
200
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400
34.
5
35
D
IND
500
1000
1500
2000
23
24
25
26
27
E
0
5
1
5
20
25
5
F
(PSS78 per 50 yr) (kg m
-3
per 50 yr) (°C per 50 yr)
Figure 3.9 | Upper 2000 dbar zonally-averaged linear trend (1950 to 2000) (colours with white contours) of salinity changes (column 1, PSS-78 per 50 yr), neutral density changes
(column 2, kg m
-3
per 50 yr), and potential temperature changes (column 3, °C per 50 yr), for the Atlantic Ocean (ATL) in row 1, Indian Ocean (IND), row 2, Pacific Ocean (PAC),
row 3, and global ocean (GLO) in row 4. Mean fields are shown as black lines (salinity: thick black contours 0.5 PSS-78, thin contours 0.25 PSS-78; neutral density: thick black
contours 1.0 kg m
-3
, thin contours 0.25 kg m
-3
; potential temperature: thick black contours 5.0°C, thin contours 2.5°C). Trends are calculated on pressure surfaces (1 dbar pressure
is approximately equal to 1 m in depth). Regions where the resolved linear trend is not significant at the 90% confidence level are stippled in grey. Salinity results are republished
from Durack and Wijffels (2010) with the unpublished temperature and density results from that study also presented.
281
3
Observations: Ocean Chapter 3
Recent studies showed that the warming of the upper ocean (Section
3.2.2) very likely affects properties of water masses in the interior, in
direct and indirect ways. Transport of SST and SSS anomalies caused
by changes in surface heat and freshwater fluxes are brought into the
ocean’s interior by contact with the surface ocean (Sections 3.2 and
3.3). Vertical and horizontal displacements of isopycnals due to surface
warming could change salinity and temperature (Section 3.3). Circu-
lation changes (Section 3.6) could also change salinity by shifting the
outcrop area of this isopycnal in regions with higher (or lower) E – P.
Properties of several deep and bottom water masses are the product of
near surface processes and significant mixing or entrainment of other
ambient water masses (Section 3.5). Changes in the properties of the
entrained or admixed water mass could dominate the observed deep
and bottom water mass changes, for instance, in the LNADW and the
AABW in the Weddell Sea.
From 1950 to 2000, it is likely that subtropical salinity maximum waters
have become more saline, while fresh intermediate waters formed at
higher latitudes have generally become fresher. In the extratropical
North Atlantic, it is very likely that the temperature, salinity, and
formation rate of the UNADW is dominated by strong decadal
variability related to NAO. It is likely that LNADW has cooled from 1955
to 2005. It is likely that the abyssal layer ventilated by AABW warmed
over much of the globe since the 1980s or 1990s respectively, and the
volume of cold AABW has been reduced over this time period.
3.6 Changes in Ocean Circulation
3.6.1 Global Observations of Ocean Circulation
Variability
The present-day ocean observing system includes global observa-
tions of velocity made at the sea surface by the Global Drifter Pro-
gram (Dohan et al., 2010), and at 1000 m depth by the Argo Program
(Freeland et al., 2010). In addition, Argo observes the geostrophic
shear between 2000 m and the sea surface. These two recently imple-
mented observing systems, if sustained, will continue to document
the large-spatial scale, long-time-scale variability of circulation in the
upper ocean. The drifter program achieved its target of 1250 drifters in
2005, and Argo its target of 3000 floats in 2007.
Historically, global measurements of ocean circulation are much spars-
er, so estimates of decadal and longer-term changes in circulation are
very limited. Since 1992, high-precision satellite altimetry has meas-
ured the time variations in sea surface height (SSH), whose horizontal
gradients are proportional to the surface geostrophic velocity. In addi-
tion, a single global top-to-bottom hydrographic survey was carried
out by the World Ocean Circulation Experiment (WOCE), mostly during
1991–1997, measuring geostrophic shear as well as velocity from mid-
depth floats and from lowered acoustic Doppler current profilers. A
subset of WOCE and pre-WOCE transects is being repeated at 5- to
10-year intervals (Hood et al., 2010).
Ocean circulation studies in relation to climate have focused on var-
iability in the wind-driven gyres (Section 3.6.2) and changes in the
meridional overturning circulations (MOCs, Sections 3.6.3 and 3.6.4)
influenced by buoyancy loss and water-mass formation as well as wind
forcing. The MOCs are responsible for much of the ocean’s capacity
to carry excess heat from the tropics to middle latitudes, and also are
important in the ocean’s sequestration of carbon. The connections
between ocean basins (Section 3.6.5) have also been subject to study
because of the significance of inter-basin exchanges in wind-driven
and thermohaline variability, and also because these can be logistically
advantageous regions for measurement (“chokepoints”). An assess-
ment is now possible of the recent mean and the changes in global
geostrophic circulation over the previous decade (Figure 3.10, and dis-
cussion in Section 3.6.2). In general, changes in the slope of SSH across
ocean basins indicate changes in the major gyres and the interior
component of MOCs. Changes occurring in high gradient regions such
as the Antarctic Circumpolar Current (ACC) may indicate shifts in the
location of those currents. In the following, the best-studied and most
significant aspects of circulation variability and change are assessed
including wind-driven circulation in the Pacific, the Atlantic and Ant-
arctic MOCs, and selected interbasin exchanges.
3.6.2 Wind-Driven Circulation Variability in the
Pacific Ocean
The Pacific covers over half of the global ocean area and its wind-
driven variability is of interest both for its consistency with wind stress
observations and for potential air–sea feedbacks that could influence
climate. Changes in Pacific Ocean circulation since the early 1990s to
the present, from the subarctic gyre to the southern ocean, observed
with satellite ocean data and in situ ocean measurements, are in good
agreement and consistent with the expected dynamical response to
observed changes in wind stress forcing.
The subarctic gyre in the North Pacific poleward of 40°N consists of the
Alaska Gyre to the east and the Western Subarctic Gyre (WSG). Since
1993, the cyclonic Alaska Gyre has intensified while decreasing in size.
The shrinking is seen in the northward shift of the North Pacific Current
Figure 3.10 | Mean steric height of the sea surface relative to 2000 decibars (black
contours at 10-cm intervals) shows the pattern of geostrophic flow for the Argo era
(2004–2012) based on Argo profile data, updated from Roemmich and Gilson (2009).
The sea surface height (SSH) trend (cm per decade, colour shading) for the period
1993–2011 is based on the AVISO altimetry “reference” product (Ducet et al., 2000).
Spatial gradients in the SSH trend, divided by the (latitude-dependant) Coriolis param-
eter, are proportional to changes in surface geostrophic velocity. For display, the mean
steric height contours and SSH trends are spatially smoothed over 5° longitude and 3°
latitude.
(cm per decade)
282
Chapter 3 Observations: Ocean
3
(NPC, the high gradient region centred about 40°N in Figure 3.10) and
has been described using the satellite altimeter, XBT/hydrography, and,
more recently, Argo profiling float data (Douglass et al., 2006; Cum-
mins and Freeland, 2007). A similar 20-year trend is detected in the
WSG, with the northern WSG in the Bering Sea having intensified while
the southern WSG south of the Aleutian Islands has weakened. These
decadal changes are attributable to strengthening and northward
expansion of the Pacific High and Aleutian Low atmospheric pressure
systems over the subarctic North Pacific Ocean (Carton et al., 2005).
The subtropical gyre in the North Pacific also expanded along its
southern boundary over the past two decades. The North Equatorial
Current (NEC) shifted southward along the 137°E meridian (Qiu and
Chen, 2012; also note the SSH increase east of the Philippines in Figure
3.10 indicating the southward shift). The NEC’s bifurcation latitude
along the Philippine coast migrated southward from a mean latitude
of 13°N in the early 1990s to 11°N in the late 2000s (Qiu and Chen,
2010). These changes are due to a recent strengthening of the Walker
circulation generating a positive wind stress curl anomaly (Tanaka et
al., 2004; Mitas and Clement, 2005). The enhanced regional sea level
rise, >10 mm yr
–1
in the western tropical North Pacific Ocean (Timmer-
mann et al., 2010, Figure 3.10), is indicative of the changes in ocean
circulation. The 20-year time-scale expansion of the North Pacific sub-
tropical gyre has high confidence owing to the good agreement seen
in satellite altimetry, subsurface ocean data and wind stress changes.
This sea level increase in the western tropical Pacific also indicates a
strengthening of the equatorward geostrophic limb of the subtropical
cells. However, the 20-year increase reversed a longer term weakening
of the subtropical cells (Feng et al., 2010), illustrating the high difficulty
of separating secular trends from multi-decadal variability.
Variability in the mid-latitude South Pacific over the past two decades
is characterized by a broad increase in SSH in the 35°S to 50°S band
and a lesser increase south of 50°S along the path of the ACC (Figure
3.10). These SSH fluctuations are induced by the intensification in the
SH westerlies (i.e., the SAM; see also Section 3.4.4), generating positive
and negative wind stress curl anomalies north and south of 50°S. In
response, the southern limb of the South Pacific subtropical gyre has
intensified in the past two decades (Cai, 2006; Qiu and Chen, 2006;
Roemmich et al., 2007) along with a southward expansion of the East
Australian Current (EAC) into the Tasman Sea (Hill et al., 2008). The
intensification in the South Pacific gyre extends to a greater depth
(>1800 m) than that in the North Pacific gyre (Roemmich and Gilson,
2009). As in the north, the 20-year changes in the South Pacific are
seen with high confidence as they occur consistently in multiple lines
of medium and high-quality data. Multiple linear regression analysis of
the 20-year Pacific SSH field (Zhang and Church, 2012) indicated that
interannual and decadal modes explain part of the circulation varia-
bility seen in SSH gradients, and once the aliasing by these modes is
removed, the SSHtrends are weaker and more spatially uniform than
in a single variable trend analysis.
The strengthening of SH westerlies is a multi-decadal signal, as seen in
SLP difference between middle and high southern latitudes from 1949
to 2009 (Gillett and Stott, 2009; also Section 3.4.4). The multi-decadal
warming in the Southern Ocean (e.g., Figure 3.1, and Gille, 2008, for
the past 50 to 70 years) is consistent with a poleward displacement of
the ACC and the southern limb of the subtropical gyres, by about 1°
of latitude per 40 years (Gille, 2008). The warming and corresponding
sea level rise signals are not confined to the South Pacific, but are seen
globally in zonal mean fields (e.g., at 40°S to 50°S in Figures 3.9 I and
3.10). Alory et al. (2007) describe the broad warming consistent with a
southward shift of the ACC in the South Indian Ocean. In the Atlantic,
a southward trend in the location of the Brazil-Malvinas confluence (at
around 39°S) is described from surface drifters and altimetry by Lump-
kin and Garzoli (2011), and in the location of the Brazil Current sep-
aration point from SST and altimetry by Goni et al. (2011). Enhanced
surface warming and poleward displacement, globally, of the western
boundary currents is described by Wu et al. (2012).
Changes in Pacific Ocean circulation over the past two decades since
1993, observed with medium to high confidence, include intensifica-
tion of the North Pacific subpolar gyre, the South Pacific subtropical
gyre, and the subtropical cells, plus expansion of the North Pacific sub-
tropical gyre and a southward shift of the ACC. It is likely that these
wind-driven changes are predominantly due to interannual-to-decadal
variability, and in the case of the subtropical cells represent reversal
of earlier multi-decadal change. Sustained time series of wind stress
forcing and ocean circulation will permit increased skill in separating
interannual and decadal variability from long-term trends (e.g., Zhang
and Church, 2012).
3.6.3 The Atlantic Meridional Overturning Circulation
The Atlantic Meridional Overturning Circulation (AMOC) consists
of an upper limb with net northward transport between the surface
and approximately 1200 m depth, and a lower limb of denser, colder,
fresher waters returning southward between 1200 m and 5000 m. The
AMOC is responsible for most of the meridional transport of heat and
carbon by the mid-latitude NH ocean and associated with the produc-
tion of about half of the global ocean’s deep waters in the northern
North Atlantic. Coupled climate models find that a slowdown of the
AMOC in the next decades is very likely, though with uncertain magni-
tude (Section 11.3.3.3). Observations of the AMOC are directed toward
detecting possible long-term changes in its amplitude, its northward
energy transport, and in the ocean’s capacity to absorb excess heat
and greenhouse gases, as well as characterizing short-term variability
and its relationship to changes in forcing.
Presently, variability in the full AMOC and meridional heat flux are
being estimated on the basis of direct observations at 26.5°N by the
RAPID/MOCHA array (Cunningham et al., 2007; Kanzow et al., 2007;
Johns et al., 2011). The array showed a mean AMOC magnitude of 18
± 1.0 Sv (±1 standard deviation of annual means) between April 2004
and April 2009, with 10-day values ranging from 3 to 32 Sv (McCarthy
et al., 2012). Earlier estimates of AMOC strength from five shipboard
expeditions over 47 years at 24°N (Bryden et al., 2005) were in the
range of variability seen by RAPID/MOCHA. For the 1-year period 1
April 2009 to 31 March 2010, the AMOC mean strength decreased
to 12.8 Sv. This decrease was manifest in a shift of southward interi-
or transport from the deep layers to the upper 1000 m. Although the
AMOC weakening in 2009/2010 was large, it subsequently rebounded
and with the large year-to-year changes no trend is detected in the
updated time-series (Figure 3.11b).
283
3
Observations: Ocean Chapter 3
2000 2002 2005 2007 2010
2012
−30
−20
−10
0
10
20
Year
Transport positive northward (Sv)
(b)
RAPID/MOCHA:
26
o
N: 17.5
±
3.8 Sv
41
o
N: 13.8
±
3.3 Sv
MOVE: 16
o
N: −20.3
±
4.8 Sv
Upper limb
Lower limb
1965 1970 1975 1980 1985 1990 1995 2000 2005 2010
20
25
30
35
40
45
50
Florida current transport (Sv)
Year
(a)
NOAA Cable
Sanford cable
NOAA dropsonde
NR dropsonde (offset by 2 Sv)
BN dropsonde (offset by 2 Sv)
Pegasus (RSMAS & NOAA)
Pegasus as dropsonde
Figure 3.11 | (a) Volume transport in Sverdrups (Sv; where 1 Sv = 10
6
m
3
s
–1
) of the Florida Current between Florida and the Bahamas, from dropsonde measurements (symbols)
and cable voltages (continuous line), extending the time-series shown in Meinen et al. (2010) (b) Atlantic Meridional Overturning Circulation (AMOC) transport estimates (Sv): 1.
RAPID/MOCHA (Rapid Climate Change programme / Meridional Ocean Circulation and Heatflux Array) at 26.5°N (red). The array monitors the top-to-bottom Atlantic wide circula-
tion, ensuring a closed mass balance across the section, and hence a direct measure of the upper and lower limbs of the AMOC. 2. 41°N (black): An index of maximum AMOC
strength from Argo float measurements in the upper 2000m only, combined with satellite altimeter data. The lower limb is not measured. 3. Meridional Overturning Variability
Experiment (MOVE) at 16°N (blue) measuring transport of North Atlantic Deep Water in the lower limb of the AMOC between 1100 m and 4800 m depth between the Caribbean
and the mid-Atlantic Ridge. This transport is thought to be representative of maximum MOC variability based on model validation experiments. The temporal resolution of the three
time series is 10 days for 16°N and 26°N and 1 month for 41°N. The data have been 3-month low-pass filtered. Means and standard deviations for the common period of 2 April
2004 to 1 April 2010 are 17.5 ± 3.8 Sv, 13.8 ± 3.3 Sv and –20.3 ± 4.8 Sv (negative indicating the southward lower limb) for 26.5°N, 41°N and 16°N respectively. The means over
this period are indicated by the horizontal line on each time series.
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Observations targeting one limb of the AMOC include Willis (2010) at
41°N combining velocities from Argo drift trajectories, Argo tempera-
ture/salinity profiles, and satellite altimeter data (Figure 3.11b). Here
the upper limb AMOC magnitude is 15.5 Sv ± 2.4 from 2002 to 2009
(Figure 3.11b). This study suggests an increase in the AMOC strength by
about 2.6 Sv from 1993 to 2010, though with low confidence because
it is based on SSH alone in the pre-Argo interval of 1993–2001. At
16°N, geostrophic array-based estimates of the southward transport
of the AMOC’s lower limb, in the depth range 1100 to 4700m, have
been made continuously since 2000 (Kanzow et al., 2008). These are
the longest continuous measurements of the southward flow of NADW
in the western basin. Whereas the period 2000 to mid-2009 suggested
a downward trend (Send et al., 2011), the updated time series (Figure
3.11b) has no apparent trend. In the South Atlantic at 35°S, estimates
of the AMOC upper limb were made using 27 high-resolution XBT tran-
sects (2002–2011) and Argo float data (Garzoli et al., 2013). The upper-
limb AMOC magnitude was 18.1 Sv ± 2.3 (1 standard deviation based
on cruise values), consistent with the NH estimates.
The continuous AMOC estimates at 16°N, 26.5°N and 41°N have
time series of length 11, 7, and 9 years respectively (Figure 3.11b). All
show a substantial variability of ~3 to 5 Sv for 3-month low-pass time
series, with a peak-to-peak interannual variability of 5Sv. The short-
ness of these time series and the relatively large interannual variability
emerging in them suggests that trend estimates be treated cautiously,
and no trends are seen at 95% confidence in any of the time series.
Continuous time series of AMOC components, longer than those of
the complete system at 26.5°N, have been obtained using moored
instrumentation. These include the inflow into the Arctic through Fram
Strait (since 1997, Schauer and Beszczynska-Möller, 2009) and through
the Barents Sea (since 1997, Ingvaldsen et al., 2004; Mauritzen et al.,
2011), dense inflows across sills between Greenland and Scotland
(since 1999 and 1995 respectively, Olsen et al., 2008; Jochumsen et
al., 2012) and North Atlantic Deep Water carried southward within the
Deep Western Boundary Current at 53°N (since 1997, Fischer et al.,
2010) and at 39°N (Line W, since 2004, Toole et al., 2011). The longest
time series of observations of ocean transport in the world (dropsonde
and cable voltage measurements in the Florida Straits), extend from
the mid-1960s to the present (Meinen et al., 2010), with small decad-
al variability of about 1 Sv and no evidence of a multi-decadal trend
(Figure 3.11a). Similarly, none of the other direct, continuous transport
estimates of single components of the AMOC exhibit long-term trends
at 95% significance.
Indirect estimates of the annual average AMOC strength and variability
can be made (Grist et al., 2009; Josey et al., 2009) from diapycnal trans-
ports driven by air–sea fluxes (NCEP-NCAR reanalysis fields from 1960
to 2007) or by inverse techniques (Lumpkin and Speer, 2007). Decadal
fluctuations of up to 2Sv are seen, but no trend. Consistent with Grist
et al. (2009), the sea level index of the strength of the AMOC, based on
several coherent western boundary tide gauge records between 39°N
and 43°N at the American coast (Bingham and Hughes, 2009) shows
no long-term trend from 1960 to 2007.
In summary, measurements of the AMOC and of circulation elements
contributing to it, at various latitudes and covering different time
periods, agree that the range of interannual variability is 5 Sv (Figure
3.11b). These estimates do not have trends, in either the subtropical or
the subpolar gyre. However, the observational record of AMOC varia-
bility is short, and there is insufficient evidence to support a finding of
change in the transport of the AMOC.
3.6.4 The Antarctic Meridional Overturning Circulation
Sinking of AABW near Antarctica supplies about half of the deep and
abyssal waters in the global ocean (Orsi et al., 1999). AABW spreads
northward as part of the global overturning circulation and ventilates
the bottom-most portions of much of the ocean. Observed widespread
warming of AABW in recent decades (Section 3.5.4) implies a con-
comitant reduction in its northward spread. Reductions of 1 to 4 Sv
in northward transports of AABW across 24°N have been estimated
by geostrophic calculations using repeat oceanographic section data
between 1981 and 2010 in the North Atlantic Ocean (Johnson et al.,
2008b; Frajka-Williams et al., 2011) and between 1985 and 2005 in
the North Pacific (Kouketsu et al., 2009). A global full-depth ocean data
assimilation study shows a reduction of northward AABW flow across
35°S of >2 Sv in the South Pacific starting around 1985 and >1 Sv in
the western South Atlantic since around 1975 (Kouketsu et al., 2011).
This reduction is consistent with the contraction in volume of AABW
(Purkey and Johnson, 2012) discussed in Section 3.5.4.
Several model studies have suggested that changes in wind stress over
the Southern Ocean (Section 3.4) may drive a change in the Southern
Ocean overturning circulation (e.g., Le Quéré et al., 2007). A recent
analysis of changes in chlorofluorocarbon (CFC) concentrations in the
Southern Ocean supports the idea that the overturning cell formed
by upwelling of deep water and sinking of intermediate waters has
slowed, but does not quantify the change in transport (Waugh et al.,
2013).
3.6.5 Water Exchange Between Ocean Basins
3.6.5.1 The Indonesian Throughflow
The transport of water from the Pacific to the Indian Ocean via the
Indonesian archipelago is the only low-latitude exchange between
oceans, and is significant because it is a fluctuating sink/source for
very warm tropical water in the two oceans. The Indonesian Through-
flow (ITF) transport has been estimated from hydrographic and XBT
transects between Australia and Indonesia, and as a synthesis of these
together with satellite altimetry, wind stress, and other data (Wunsch,
2010), and from moorings in the principal Indonesian passages. The
most comprehensive observations were obtained in 2004–2006 in
three passages by the INSTANT mooring array (Sprintall et al., 2009),
and show a westward transport of 15.0 (±4) Sv. For the main pas-
sage, Makassar Strait, Susanto et al. (2012) find 13.3 (±3.6) Sv in the
period 2004–2009, with small year-to-year differences. On a longer
time scale, the Wunsch (2010) estimate for 1992–2007 was 11.5 Sv
(±2.4) westward, and thus consistent with INSTANT. Wainwright et al.
(2008) analyzed data between Australia and Indonesiabeginning in
the early 1950s, and found a change in the slope of the thermocline
for data before and after 1976, indicating a decrease in geostrophic
transport by 23%, consistent with a weakening of the tradewinds (e.g.,
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Observations: Ocean Chapter 3
Vecchi et al. (2006), who described a downward trend in the Walker cir-
culation since the late 19th century). Other transport estimates based
on the IX1 transect show correlation with ENSO variability (Potemra
and Schneider, 2007) and no significant trend for the period since 1984
having continuous sampling along IX1 (Sprintall et al., 2002). Overall,
the limited evidence provides low confidence that a trend in ITF trans-
port has been observed.
3.6.5.2 The Antarctic Circumpolar Current
There is medium confidence that the westerly winds in the Southern
Ocean have increased since the early 1980s (Section 3.4.4), associated
with a positive trend in the SAM (Marshall, 2003); also see Sections
3.4.4 and 3.6.3). Although a few observational studies have found
evidence for correlation between SAM and ACC transport on subsea-
sonal to interannual scales (e.g., Hughes et al., 2003; Meredith et al.,
2004), there is no significant observational evidence of an increase in
ACC transport associated with the multi-decadal trend in wind forcing
over the Southern Ocean. Repeat hydrographic sections spread une-
venly over 35 years in Drake Passage (e.g., Cunningham et al., 2003;
Koshlyakov et al., 2007, 2011; Gladyshev et al., 2008), south of Africa
(Swart et al., 2008) and south of Australia (Rintoul et al., 2002) reveal
moderate variability but no significant trends in these sparse and dis-
continuous records. A comparison of recent Argo data and a long-term
climatology showed that the slope of density surfaces (hence baroclinic
transport) associated with the ACC had not changed in recent decades
(Böning et al., 2008). Eddy-resolving models suggest the ACC trans-
port is relatively insensitive to trends in wind forcing, consistent with
the ACC being in an “eddy-saturated” state where increases in wind
forcing are compensated by changes in the eddy field (Hallberg and
Gnanadesikan, 2006; Farneti et al., 2010; Spence et al., 2010). While
there is limited evidence for (or against) multi-decadal changes in
transport of the ACC, observations of changes in temperature, salinity
and SSH indicate the current system has shifted poleward (medium
confidence) (Böning et al., 2008; Gille, 2008; Morrow et al., 2008; Soko-
lov and Rintoul, 2009; Kazmin, 2012).
3.6.5.3 North Atlantic/Nordic Seas Exchange
There is no observational evidence of changes during the past two
decades in the flow across the Greenland–Scotland Ridge, which con-
nects the North Atlantic with the Norwegian and Greenland Seas.
Direct current measurements since the mid-1990s have not shown
any significant trends in volume transport for any of the three inflow
branches (Østerhus et al., 2005; Hansen et al., 2010; Mauritzen et al.,
2011; Jónsson and Valdimarsson, 2012).
The two primary pathways for the deep southward overflows across
the Greenland–Scotland Ridge are the Denmark Strait and Faroe
Bank Channel. Moored measurements of the Denmark Strait overflow
demonstrate significant interannual transport variations (Macrander
et al., 2005; Jochumsen et al., 2012), but the time series is not long
enough to detect a multi-decadal trend. Similarly, a 10-year time series
of moored measurements in the Faroe Bank channel (Olsen et al.,
2008) does not show a trend in transport.
3.6.6 Conclusions
Recent observations have greatly increased the knowledge of the
amplitude of variability in major ocean circulation systems on time
scales from years to decades. It is very likely that the subtropical gyres
in the North Pacific and South Pacific have expanded and strength-
ened since 1993, but it is about as likely as not that this reflects a
decadal oscillation linked to changes in wind forcing, including chang-
es in winds associated with the modes of climate variability. There is
no evidence for a long-term trend in the AMOC amplitude, based on
a decade of continuous observations plus several decades of sparse
hydrographic transects, or in the longer records of components of the
AMOC such as the Florida Current (since 1965), although there are
large interannual fluctuations. Nor is there evidence of a trend in the
transports of the ITF (over about 20 years), the ACC (about 30 years
sparsely sampled), or between the Atlantic and Nordic Seas (about 20
years). Given the short duration of direct measurements of ocean cir-
culation, we have very low confidence that multi-decadal trends can
be separated from decadal variability.
3.7 Sea Level Change, Including Extremes
3.7.1 Introduction and Overview of Sea Level
Measurements
Sea level varies as the ocean warms or cools, as water is transferred
between the ocean and continents, between the ocean and ice sheets,
and as water is redistributed within the ocean due to the tides and
changes in the oceanic and atmospheric circulation. Sea level can rise
or fall on time scales ranging from hours to centuries, spatial scales
from <1 km to global, and with height changes from a few millimeters
to a meter or more (due to tides). Sea level integrates and reflects
multiple climatic and dynamical signals. Measurements of sea level are
the longest-running ocean observation system. This section assesses
interannual and longer variations in non-tidal sea level from the instru-
mented period (late 18th century to the present). Sections 4.3.3 and
4.4.2 assess contributions of glaciers and ice sheets to sea level, Sec-
tion 5.6 assess reconstructions of sea level from the geological record,
Section 10.4.3 assesses detection and attribution of human influences
on sea level change, and Chapter 13 synthesizes results and assesses
projections of sea level change.
The sea level observing system has evolved over time. There are inter-
mittent records of sea level at four sites in Northern Europe starting
in the 1700s. By the late 1800s, there were more tide gauges being
operated in Northern Europe, on both North American coasts, and in
Australia and New Zealand in the SH (Appendix 3.A). Tide gauges
began to be placed on islands far from continental coasts starting in
the early 20th century, but a majority of deep-ocean islands did not
have an operating tide gauge suitable for climate studies until the
early 1970s.
Tide gauge records measure the combined effect of ocean volume
change and vertical land motion (VLM). For detecting climate related
variability of the ocean volume, the VLM signal must be removed. One
component that can be accounted for to a certain extent is the VLM
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3
associated with glacial isostatic adjustment (GIA) (Peltier, 2001). In
some areas, however, VLM from tectonic activity, groundwater mining,
or hydrocarbon extraction is greater than GIA (e.g., Wöppelmann et
al., 2009; King et al., 2012); these effects can be reduced by selecting
gauges with no known tectonic or subsidence issues (e.g., Douglas,
2001) or by selecting gauges where GIA models have small differences
(Spada and Galassi, 2012). More recently, Global Positioning System
(GPS) receivers have been installed at tide gauge sites to measure VLM
as directly as possible (e.g., Wöppelmann et al., 2009; King et al., 2012).
However, these measurements of VLM are only available since the late
1990s at the earliest, and either have to be extrapolated into the past to
apply to older records, or used to identify sites without extensive VLM.
Satellite radar altimeters in the 1970s and 1980s made the first nearly
global observations of sea level, but these early measurements were
highly uncertain and of short duration. The first precise record began
with the launch of TOPEX/Poseidon (T/P) in 1992. This satellite and its
successors (Jason-1, Jason-2) have provided continuous measurements
of sea level variability at 10-day intervals between approximately ±66°
latitude. Additional altimeters in different orbits (ERS-1, ERS-2, Envi-
sat, Geosat Follow-on) have allowed for measurements up to ±82°
latitude and at different temporal sampling (3 to 35 days), although
these measurements are not as accurate as those from the T/P and
Jason satellites. Unlike tide gauges, altimetry measures sea level rela-
tive to a geodetic reference frame (classically a reference ellipsoid that
coincides with the mean shape of the Earth, defined within a globally
realized terrestrial reference frame) and thus will not be affected by
VLM, although a small correction that depends on the area covered by
the satellite (~0.3 mm yr
–1
) must be added to account for the change in
location of the ocean bottom due to GIA relative to the reference frame
of the satellite (Peltier, 2001; see also Section 13.1.2).
Tide gauges and satellite altimetry measure the combined effect of
ocean warming and mass changes on ocean volume. Although var-
iations in the density related to upper-ocean salinity changes cause
regional changes in sea level, when globally averaged their effect on
sea level rise is an order of magnitude or more smaller than thermal
effects (Lowe and Gregory, 2006). The thermal contribution to sea level
can be calculated from in situ temperature measurements (Section
3.2). It has only been possible to directly measure the mass compo-
nent of sea level since the launch of the Gravity Recovery and Climate
Experiment (GRACE) in 2002 (Chambers et al., 2004). Before that, esti-
mates were based either on estimates of glacier and ice sheet mass
losses or using residuals between sea level measured by altimetry or
tide gauges and estimates of the thermosteric component (e.g., Willis
et al., 2004; Domingues et al., 2008), which allowed for the estima-
tion of seasonal and interannual variations as well. GIA also causes a
gravitational signal in GRACE data that must be removed in order to
determine present-day mass changes; this correction is of the same
order of magnitude as the expected trend and is still uncertain at the
30% level (Chambers et al., 2010).
3.7.2 Trends in Global Mean Sea Level and Components
Tide gauges with the longest nearly continuous records of sea
level show increasing sea level over the 20th century (Figure 3.12;
Woodworth et al., 2009; Mitchum et al., 2010). There are, however,
significant interannual and decadal-scale fluctuations about the aver-
age rate of sea level rise in all records. Different approaches have been
used to compute the mean rate of 20th century global mean sea level
(GMSL) rise from the available tide gauge data: computing average
rates from only very long, nearly continuous records (Douglas, 2001;
Holgate, 2007); using more numerous but shorter records and filters to
separate nonlinear trends from decadal-scale quasi-periodic variability
(Jevrejeva et al., 2006, 2008); neural network methods (Wenzel and
Schroeter, 2010); computing regional sea level for specific basins then
Figure 3.12 | 3-year running mean sea level anomalies (in millimeters) relative to
1900–1905 from long tide gauge records representing each ocean basin from the
Permanent Service for Mean Sea Level (PSMSL) (http://www.psmsl.org), obtained May
2011. Data have been corrected for Glacial Isostatic Adjustment (GIA) (Peltier, 2004),
using values available from http://www.psmsl.org/train_and_info/geo_signals/gia/pel-
tier/. Error bars reflect the 5 to 95% confidence interval, based on the residual monthly
variability about the 3-year running mean.
-100
-50
0
50
100
150
200
1880 1900 1920 1940 1960 1980 2000
MSL anomaly (mm)
Year
New York City, USA
Newlyn, UK
-100
-50
0
50
100
150
200
1880 1900 1920 1940 1960 1980 2000
MSL anomaly (mm)
Year
San Francisco, USA
Sydney, Australia
-100
-50
0
50
100
150
200
1880 1900 1920 1940 1960 1980 2000
MSL anomaly (mm)
Year
Mumbai, India
Fremantle, Australia
(a)
(b)
(c)
287
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Observations: Ocean Chapter 3
averaging (Jevrejeva et al., 2006, 2008; Merrifield et al., 2009; Wöp-
pelmann et al., 2009); or projecting tide gauge records onto empirical
orthogonal functions (EOFs) computed from modern altimetry (Church
et al., 2004; Church and White, 2011; Ray and Douglas, 2011) or EOFs
from ocean models (Llovel et al., 2009; Meyssignac et al., 2012). Dif-
ferent approaches show very similar long-term trends, but noticeably
different interannual and decadal-scale variability (Figure 3.13a). Only
the time series fromChurch and White(2011) extends to 2010, so it is
used in the assessment of rates of sea level rise.The rate from 1901 to
2010 is 1.7 [1.5 to 1.9] mm yr
–1
(Table 3.1), which is unchanged from
the value in AR4. Rates computed using alternative approaches over
the longest common interval (1900–2003) agree with this estimate
within the uncertainty.
Since AR4, significant progress has been made in quantifying the uncer-
tainty in GMSL associated with unknown VLM and uncertainty in GIA
models. Differences between rates of GMSL rise computed with and
without VLM from GPS are smaller than the estimated uncertainties
(Merrifield et al., 2009; Wöppelmann et al., 2009). Use of different GIA
models to correct tide gauge measurements results in differences less
than 0.2 mm yr
–1
(one standard error), and rates of GMSL rise com-
puted from uncorrected tide gauges differ from rates computed from
GIA-corrected gauges by only 0.4 mm yr
–1
(Spada and Galassi, 2012),
again within uncertainty estimates. This agreement gives increased
confidence that the 20th century rate of GMSL rise is not biased high
due to unmodeled VLM at the gauges.
Satellite altimetry can resolve interannual fluctuations in GMSL better
than tide gauge records because less temporal smoothing is required
(Figure 3.13b). It is clear that deviations from the long-term trend can
exist for periods of several years, especially during El Niño (e.g., 1997–
1998) and La Niña (e.g., 2011) events (Nerem et al., 1999; Boening et
al., 2012; Cazenave et al., 2012). The rate of GMSL rise from 1993–
2010 is 3.2 [2.8 to 3.6] mm yr
–1
based on the average of altimeter time
series published by multiple groups (Ablain et al., 2009; Beckley et al.,
2010; Leuliette and Scharroo, 2010; Nerem et al., 2010; Church and
White, 2011; Masters et al., 2012, Figure 3.13). As noted in AR4, this
rate continues to be statistically higher than that for the 20th century
-10
0
10
20
30
40
50
60
70
1992
1994 1996 1998 2000 2002 2004 2006 2008 2010 2012
Year
Tide gauge
Altimeter
-20
0
20
40
60
80
100
1970
1975 1980 1985 1990 1995 2000 2005 2010
Year
Sea level
Thermosteric component
-10
-5
0
5
10
15
2005
2006 2007 2008 2009 2010 2011 2012
Year
Mass (GRACE) + Steric (ARGO)
Sea level (Altimeter)
(a)
(b)
(d)
(c)
GMSL anomaly (mm)
GMSL anomaly (mm)
GMSL anomaly (mm)
-100
-50
0
50
100
150
200
1880 1900 1920 1940 1960 1980 2000
Church & White, 2011
Jevrejeva et al., 2008
Ray & Douglas, 2011
GMSL anomaly (mm)
Year
Figure 3.13 | Global mean sea level anomalies (in mm) from the different measuring systems as they have evolved in time, plotted relative to 5-year mean values that start at (a)
1900, (b) 1993, (c) 1970 and (d) 2005. (a) Yearly average GMSL reconstructed from tide gauges (1900–2010) by three different approaches (Jevrejeva et al., 2008; Church and
White, 2011; Ray and Douglas, 2011). (b) GMSL (1993–2010) from tide gauges and altimetry (Nerem et al., 2010) with seasonal variations removed and smoothed with a 60-day
running mean. (c) GMSL (1970–2010) from tide gauges along with the thermosteric component to 700 m (3-year running mean) estimated from in situ temperature profiles
(updated from Domingues et al., 2008). (d) The GMSL (nonseasonal) from altimetry and that computed from the mass component (GRACE) and steric component (Argo) from 2005
to 2010 (Leuliette and Willis, 2011), all with a 3-month running mean filter. All uncertainty bars are one standard error as reported by the authors. The thermosteric component is
just a portion of total sea level, and is not expected to agree with total sea level.
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3
(Table 3.1). There is high confidence that this change is real and not
an artefact of the different sampling or change in instrumentation,
as the trends estimated over the same period from tide gauges and
altimetry are consistent. Although the rate of GMSL rise has a slightly
lower trend between 2005 and 2010 (Nerem et al., 2010), this variation
is consistent with earlier interannual fluctuations in the record (e.g.,
1993–1997), mostly attributable to El Niño/La Niña cycles (Box 9.2).
At least 15 years of data are required to reduce the impact of interan-
nual variations associated with El Niño or La Niña on estimated trends
(Nerem et al., 1999).
Since AR4, estimates of both the thermosteric component and mass
component of GMSL rise have improved, although estimates of the
mass component are possible only since the start of the GRACE meas-
urements in 2002. After correcting for biases in older XBT data [3.2],
the rate of thermosteric sea level rise in the upper 700 m since 1971
is 50% higher than estimates used for AR4 (Domingues et al., 2008;
Wijffels et al., 2008). Because of much sparser upper ocean measure-
ments before 1971, we estimate the trend only since then (Section
3.2). The warming of the upper 700 m from 1971 to 2010 caused an
estimated mean thermosteric rate of rise of 0.6 [0.4 to 0.8] mm yr
–1
(90% confidence), which is 30% of the observed rate of GMSL rise
for the same period (Table 3.1; Figure 3.13c). Although still a short
record, more numerous, better distributed, and higher quality profile
measurements from the Argo program are now being used to estimate
the steric component for the upper 700 m as well as for the upper
2000 m (Domingues et al., 2008; Willis et al., 2008, 2010; Cazenave et
al., 2009; Leuliette and Miller, 2009; Leuliette and Willis, 2011; Llovel
et al., 2011; von Schuckmann and Le Traon, 2011; Levitus et al., 2012).
However, these data have been shown to be best suited for global
analyses after 2005 owing to a combination of interannual variabil-
ity and large biases when using data before 2005 owing to sparser
sampling (Leuliette and Miller, 2009; von Schuckmann and Le Traon,
2011). Comparison of sparse but accurate temperature measurements
from the World Ocean Circulation Experiment in the 1990s with Argo
data from 2006 to 2008 also indicates a significant rise in global ther-
mosteric sea level, although the estimate is uncertain owing to rela-
tively sparse 1990s sampling (Freeland and Gilbert, 2009).
Observations of the contribution to sea level rise from warming below
700 m are still uncertain due to limited historical data, especially in the
Southern Ocean (Section 3.2). Before Argo, they are based on 5-year
averages to 2000 m depth (Levitus et al., 2012). From 1971 to 2010,
the estimated trend for the contribution between 700 m and 2000 m
is 0.1 [0 to 0.2] mm yr
–1
(Table 3.1; Levitus et al., 2012). To measure
the contribution of warming below 2000 m, much sparser but very
accurate temperature profiles along repeat hydrographic sections are
utilized (Purkey and Johnson, 2010; Kouketsu et al., 2011). The studies
have found a significant warming trend between 1000 and 4000 m
within and south of the Sub-Antarctic Front (Figure 3.3). The estimated
total contribution of warming below 2000 m to global mean sea level
rise between about 1992 and 2005 is 0.1 [0.0 to 0.2] mm yr
–1
(95%
confidence as reported by authors; Purkey and Johnson, 2010).
Detection of the mass component of sea level from the GRACE mis-
sion was not assessed in AR4, as the record was too short and there
was still considerable uncertainty in the measurements and corrections
required. Considerable progress has been made since AR4, and the
mass component of sea level measured by GRACE has been increas-
ing at a rate between 1 and 2 mm yr
–1
since 2002 (Willis et al., 2008,
2010; Cazenave et al., 2009; Leuliette and Miller, 2009; Chambers et
al., 2010; Llovel et al., 2010; Leuliette and Willis, 2011). Differences
between studies are due partially to the time periods used to com-
pute trends, as there are significant interannual variations in the mass
component of GMSL (Willis et al., 2008; Chambers et al., 2010; Llovel
et al., 2010; Boening et al., 2012), but also to substantial differences in
GIA corrections applied, of order 1 mm yr
–1
. Recent evaluations of the
GIA correction have found explanations for the difference (Chambers
et al., 2010; Peltier et al., 2012), but uncertainty of 0.3 mm yr
–1
is still
probable. Measurements of sea level from altimetry and the sum of
observed steric and mass components are also consistent at monthly
scales during the time period when Argo data have global distribu-
tion (Figure 3.13d), which gives high confidence that the current ocean
observing system is capable of resolving the rate of sea level rise and
its components.
3.7.3 Regional Distribution of Sea Level Change
Large-scale spatial patterns of sea level change are known to high
precision only since 1993, when satellite altimetry became available
(Figure 3.10). These data have shown a persistent pattern of change
since the early 1990s in the Pacific, with rates of rise in the Warm Pool
of the western Pacific up to three times larger than those for GMSL,
while rates over much of the eastern Pacific are near zero or nega-
tive (Beckley et al., 2010). The increasing sea level in the Warm Pool
started shortly before the launch of TOPEX/Poseidon (Merrifield, 2011),
and is caused by an intensification of the trade winds (Merrifield and
Maltrud, 2011) since the late 1980s that may be related to the Pacific
Decadal Oscillation (PDO) (Merrifield et al., 2012; Zhang and Church,
2012). The lower rate of sea level rise since 1993 along the western
coast of the United States has also been attributed to changes in the
wind stress curl over the North Pacific associated with the PDO (Bro-
mirski et al., 2011). While global maps can be created using EOF anal-
ysis (e.g., Church et al., 2004; Llovel et al., 2009), pre-1993 results are
still uncertain, as the method assumes that the EOFs since 1993 are
capable of representing the patterns in previous decades, and results
may be biased in the middle of the ocean where there are no tide
gauges to constrain the estimate (Ray and Douglas, 2011). Several
studies have examined individual long tide gauge records in the North
Atlantic and found coherent decadal-scale fluctuations along both the
USA east coast (Sturges and Hong, 1995; Hong et al., 2000; Miller and
Douglas, 2007), the European coast (Woodworth et al., 2010; Sturges
and Douglas, 2011; Calafat et al., 2012), and the marginal seas in the
western North Pacific (Marcos et al., 2012), all related to natural cli-
mate variability.
There is still considerable uncertainty on how long large-scale pat-
terns of regional sea level change can persist, especially in the Pacif-
ic where the majority of tide gauge records are less than 40 years
long. Based on analyses of the longest records in the Atlantic, Indian
and Pacific Oceans (including the available gauges in the Southern
Ocean) there are significant multi-decadal variations in regional sea
level (Holgate, 2007; Woodworth et al., 2009, 2011; Mitchum et al.,
2010; Chambers et al., 2012). Hence local rates of sea level rise can
289
3
Observations: Ocean Chapter 3
be considerably higher or lower than the global mean rate for periods
of a decade or more.
The preceding discussion of regional sea level trends has focused on
effects that appear to be related to regional ocean volume change, and
not those due to vertical land motion. As discussed in Section 3.7.1,
vertical land motion can dramatically affect local sea level change.
Some extreme examples of vertical land motion are in Neah Bay,
Washington, where the signal is +3.8 mm yr
–1
(uplift from tectonic
activity); Galveston, Texas, where the value is –5.9 mm yr
–1
(subsid-
ence from groundwater mining); and Nedre Gavle, Sweden where the
value is +7.1 mm yr
–1
(uplift from GIA), all computed from nearby GPS
receivers (Wöppelmann et al., 2009). These areas will all have long-
term rates of sea level rise that are significantly higher or lower than
those due to ocean volume change alone, but as these rates are not
related to climate change, they are not discussed here.
3.7.4 Assessment of Evidence for Accelerations
in Sea Level Rise
AR4 concluded that there was “high confidence that the rate of global
sea level rise increased from the 19th to the 20th century” but could
not be certain as to whether the higher rate since 1993 was reflective of
decadal variability or a further increase in the longer-term trend. Since
AR4, there has been considerable effort to quantify the level of decadal
and multi-decadal variability and to detect acceleration in GMSL and
mean sea level at individual tide gauges. It has been clear for some
time that there was a significant increase in the rate of sea level rise
in the four oldest records from Northern Europe starting in the early
to mid-19th century (Ekman, 1988; Woodworth, 1990, 1999; Mitchum
et al., 2010). Estimates of the change in the rate have been computed,
either by comparing trends over 100-year intervals for the Stockholm
site (Ekman, 1988; Woodworth, 1990), or by fitting a quadratic term to
all the long records starting before 1850 (Woodworth, 1990, 1999). The
results are consistent and indicate a significant acceleration that start-
ed in the early to mid-19th century (Woodworth, 1990, 1999), although
some have argued it may have started in the late 1700s (Jevrejeva et
al., 2008). The increase in the rate of sea level rise at Stockholm (the
longest record that extends past 1900) has been based on differenc-
ing 100-year trends from 1774–1884 and 1885–1985. The estimated
change is 1.0 [0.7 to 1.3] mm yr
–1
per century (1 standard error, as cal-
culated by Woodworth, 1990). Although sites in other ocean basins do
show an increased trend after 1860 (e.g., Figure 3.12), it is impossible
to detect a change in the early to mid-1800s in other parts of the ocean
using tide gauge data alone, as there are no observations.
Numerous studies have attempted to quantify if a detectable accelera-
tion has continued into the 20th century, typically by fitting a quadratic
to data at individual tide gauges (Woodworth, 1990; Woodworth et
al., 2009, 2011; Houston and Dean, 2011; Watson, 2011) as well as to
reconstructed time series of GMSL (Church and White, 2006; Jevrejeva
et al., 2008; Church and White, 2011; Rahmstorf and Vermeer, 2011), or
by examining differences in long-term rates computed at different tide
gauges (Sallenger et al., 2012). Woodworth et al. (2011) find significant
quadratic terms at the sites that begin before 1860 (all in the NH).
Other authors using more numerous but significantly shorter records
have found either insignificant or small negative quadratic terms in sea
level around the United States and Australia since 1920 (Houston and
Dean, 2011; Watson, 2011), or large positive quadratic values since
1950 along the U.S. east coast (Sallenger et al., 2012). However, fitting
a quadratic term to tide gauge data after 1920 results in highly varia-
ble, insignificant quadratic terms (Rahmstorf and Vermeer, 2011), and
so only studies that use data before 1920 and that extend until 2000 or
beyond are suitable for evaluating long-term acceleration of sea level.
A long time scale is needed because significant multi-decadal varia-
bility appears in numerous tide gauge records during the 20th century
(Holgate, 2007; Woodworth et al., 2009, 2011; Mitchum et al., 2010;
Chambers et al., 2012). The multi-decadal variability is marked by an
increasing trend starting in 1910–1920, a downward trend (i.e., level-
ing of sea level if a long-term trend is not removed) starting around
1950, and an increasing trend starting around 1980. The pattern can
be seen in New York, Mumbai and Fremantle records, for instance
(Figure 3.12), as well as 14 other gauges representing all ocean basins
(Chambers et al., 2012), and in all reconstructions (Figure 3.14). It is
also seen in an analysis of upper 400 m temperature (Gouretski et
al., 2012; Section 3.3.2). Although the calculations of 18-year rates of
GMSL rise based on the different reconstruction methods disagree by
as much as 2 mm yr
–1
before 1950 and on details of the variability
(Figure 3.14), all do indicate 18-year trends that were significantly
higher than the 20th century average at certain times (1920–1950,
1990–present) and lower at other periods (1910–1920, 1955–1980),
likely related to multi-decadal variability. Several studies have suggest-
ed these variations may be linked to climate fluctuations like the Atlan-
tic Multi-decadal Oscillation (AMO) and/or Pacific Decadal Oscillation
(PDO, Box 2.5) (Holgate, 2007; Jevrejeva et al., 2008; Chambers et al.,
2012), but these results are not conclusive.
While technically correct that these multi-decadal changes represent
acceleration/deceleration of sea level, they should not be interpreted
as change in the longer-term rate of sea level rise, as a time series
longer than the variability is required to detect those trends. Using data
18-year GMSL trends (mm yr
-1
)
Year
-1
0
1
2
3
4
5
1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000
Church & White
Jevrejeva et al.
Ray & Douglas
Altimeter
Figure 3.14 | 18-year trends of GMSL rise estimated at 1-year intervals. The time is
the start date of the 18-year period, and the shading represents the 90% confidence.
The estimate from satellite altimetry is also given, with the 90% confidence given as
an error bar. Uncertainty is estimated by the variance of the residuals about the fit, and
accounts for serial correlation in the residuals as quantified by the lag-1 autocorrelation.
290
Chapter 3 Observations: Ocean
3
extending from 1900 to after 2000, the quadratic term computed from
both individual tide gauge records and GMSL reconstructions is signif-
icantly positive (Jevrejeva et al., 2008; Church and White, 2011; Rahm-
storf and Vermeer, 2011; Woodworth et al., 2011). Church and White
(2006) report that the estimated acceleration term in GMSL (twice the
quadratic parameter) is 0.009 [0.006 to 0.012] mm yr
–2
(1 standard
deviation) from 1880 to 2009, which is consistent with the other pub-
lished estimates (e.g., Jevrejeva et al., 2008; Woodworth et al., 2009)
that use records longer than 100 years. Chambers et al. (2012) find that
modelling a period near 60 years removes much of the multi-decadal
variability of the 20th century in the tide gauge reconstruction time
series. When a 60-year oscillation is modeled along with an accelera-
tion term, the estimated acceleration in GMSL since 1900 ranges from:
0.000 [–0.002 to 0.002] mm yr
–2
in the Ray and Douglas (2011) record,
0.013 [0.007 to 0.019] mm yr
–2
in the Jevrejeva et al. (2008) record,
and 0.012 [0.009 to 0.015] mm yr
–2
in the Church and White (2011)
record. Thus, while there is more disagreement on the value of a 20th
century acceleration in GMSL when accounting for multi-decadal fluc-
tuations, two out of three records still indicate a significant positive
value. The trend in GMSL observed since 1993, however, is not signif-
icantly larger than the estimate of 18-year trends in previous decades
(e.g., 1920–1950).
3.7.5 Changes in Extreme Sea Level
Aside from non-climatic events such as tsunamis, extremes in sea level
(i.e., coastal flooding, storm surge, high water events, etc.) tend to be
caused by large storms, especially when they occur at times of high
tide. However, any low-pressure system offshore with associated high
winds can cause a coastal flooding event depending on the duration
and direction of the winds. Evaluation of changes in frequency and
intensity of storms have been treated in Sections 2.6.3 and 2.6.4, as
well as SREX Chapter 3 (Section 3.5.2). The main conclusions from both
are that there is low confidence of any trend or long term change in
tropical or extratropic storm frequency or intensity in any ocean basin,
although there is robust evidence for an increase in the most intense
tropical cyclones in the North Atlantic basin since the 1970s. The mag-
nitude and frequency of extreme events can still increase without a
change in storm intensity, however, if the mean water level is also
increasing. AR4 concluded that the highest water levels have been
increasing since the 1950s in most regions of the world, caused mainly
by increasing mean sea level. Studies published since AR4 continue to
support this conclusion, although higher regional extremes are also
caused by large interannual and multi-decadal variations in sea level
associated with climate fluctuations such as ENSO, the North Atlantic
Oscillation and the Atlantic Multi-decadal Oscillation, among others
(e.g., Abeysirigunawardena and Walker, 2008; Haigh et al., 2010;
Menéndez and Woodworth, 2010; Park et al., 2011).
Global analyses of the changes in extreme sea level are limited, and
most reports are based on analysis of regional data (see Lowe et al.,
2010 for a review). Estimates of changes in extremes rely either on the
analysis of local tide gauge data, or on multi-decadal hindcasts of a
dynamical model (WASA-Group, 1998). Most analyses have focused
on specific regions and find that extreme values have been increas-
ing since the 1950s, using various statistical measures such as annual
maximum surge, annual maximum surge-at-high-water, monthly mean
high water level, changes in number of high storm surge events, or
changes in 99th percentile events (e.g., Church et al., 2006; D’Onofrio
et al., 2008; Marcos et al., 2009; Haigh et al., 2010; Letetrel et al., 2010;
Tsimplis and Shaw, 2010; Vilibic and Sepic, 2010; Grinsted et al., 2012).
A global analysis of tide gauge records has been performed for data
from the 1970s onwards when the global data sampling has been
robust and finds that the magnitude of extreme sea level events has
increased in all regions studied since that time (Woodworth and Black-
man, 2004; Menéndez and Woodworth, 2010; Woodworth et al., 2011).
The height of a 50-year flood event has increased anywhere from 2
to more than 10 cm per decade since 1970 (Figure 3.15a), although
some areas have seen a negative rate because vertical land motion is
much larger than the rate of mean sea level rise. However, when the
annual median height at each gauge is removed to reduce the effect
of local mean sea level rise, interannual and decadal fluctuations, and
vertical land motion, the rate of extreme sea level change drops in
49% of the gauges to below significance (Figure 3.15b), while at 45%
it fell to less than 5 mm yr
–1
. Only 6% of tide gauge records evaluated
had a change in the amplitude of more than 5 mm yr
–1
after removing
mean sea level variations, mainly in the southeast United States, the
western Pacific, Southeast Asia and a few locations in Northern Europe.
The higher rates in the southeastern United States have been linked to
larger storm surge events unconnected to global sea level rise (Grin-
sted et al., 2012).
−12
−10 −8 −6 −4
-2 0 2 4 6 8 10 12
(cm per decade)
(a)
(b)
Figure 3.15 | Estimated trends (cm per decade) in the height of a 50-year event in
extreme sea level from (a) total elevation and (b) total elevation after removal of annual
medians. Only trends significant at the 95% confidence level are shown. (Data are from
Menéndez and Woodworth, 2010.)
291
3
Observations: Ocean Chapter 3
3.7.6 Conclusions
It is virtually certain that globally averaged sea level has risen over the
20th century, with a very likely mean rate between 1900 and 2010 of
1.7 [1.5 to 1.9] mm yr
–1
and 3.2 [2.8 and 3.6] mm yr
–1
between 1993
and 2010. This assessment is based on high agreement among multi-
ple studies using different methods, and from independent observing
systems (tide gauges and altimetry) since 1993. It is likely that a rate
comparable to that since 1993 occurred between 1920 and 1950, pos-
sibly due to a multi-decadal climate variation, as individual tide gauges
around the world and all reconstructions of GMSL show increased
rates of sea level rise during this period. Although local vertical land
motion can cause even larger rates of sea level rise (or fall) relative to
the coastline, it is very likely that this does not affect the estimates of
the global average rate, based on multiple estimations of the average
with and without VLM corrections.
It is virtually certain that interannual and decadal changes in the
large-scale winds and ocean circulation can cause significantly higher
or lower rates over shorter periods at individual locations, as this has
been observed in tide gauge records around the world. Warming of the
upper 700 m of the ocean has very likely contributed an average of 0.6
[0.4 to 0.8] mm yr
–1
of sea level change since 1971. Warming between
700 m and 2000 m has likely been contributing an additional 0.1 mm
yr
–1
[0 to 0.2] of sea level rise since 1971, and warming below 2000
m likely has been contributing another 0.1 [0.0 to 0.2] mm yr
–1
of sea
level rise since the early 1990s.
It is very likely that the rate of mean sea level rise along Northern
European coastlines has accelerated since the early 1800s and that this
has continued through the 20th century, as the increased rate since
1875 has been observed in multiple long tide gauge records and by
different groups using different analysis techniques. It is likely that sea
level rise throughout the NH has also accelerated since 1850, as this is
also observed in a smaller number of gauges along the coast of North
America. Two of the three time series based on reconstructing GMSL
from tide gauge data back to 1900 or earlier indicate a significant
positive acceleration, while one does not. The range is –0.002 to 0.019
mm yr
–2
, so it is likely that GMSL has accelerated since 1900. Finally, it
is likely that extreme sea levels have increased since 1970, largely as a
result of the rise in mean sea level.
3.8 Ocean Biogeochemical Changes, Including
Anthropogenic Ocean Acidification
The oceans can store large amounts of CO
2
. The reservoir of inorganic
carbon in the ocean is roughly 50 times that of the atmosphere (Sabine
et al., 2004). Therefore even small changes in the ocean reservoir can
have an impact on the atmospheric concentration of CO
2
. The ocean
Quantity Period
Trend
(mm yr
–1
)
Source Resolution
1901–2010 1.7 [1.5 to 1.9] Tide Gauge Reconstruction (Church and White, 2011) Yearly
1901–1990 1.5 [1.3 to 1.7] Tide Gauge Reconstruction (Church and White, 2011) Yearly
GMSL
1971–2010 2.0 [1.7 to 2.3] Tide Gauge Reconstruction (Church and White, 2011) Yearly
1993–2010 2.8 [2.3 to 3.3] Tide Gauge Reconstruction (Church and White, 2011) Yearly
1993–2010 3.2 [2.8 to 3.6]
a
Altimetry (Nerem et al., 2010) time-series 10-Day
Thermosteric Component
(upper 700 m)
1971–2010 0.6 [0.4 to 0.8] XBT Reconstruction (updated from Domingues et al., 2008) 3-Year running means
1993–2010 0.8 [0.5 to 1.1] XBT Reconstruction (updated from Domingues et al., 2008) 3-Year running means
Thermosteric Component
(700 to 2000 m)
1971–2010 0.1 [0 to 0.2] Objective mapping of historical temperature data (Levitus et al., 2012) 5-Year averages
1993–2010 0.2 [0.1 to 0.3] Objective mapping of historical temperature data (Levitus et al., 2012) 5-Year averages
Thermosteric Component
(below 2000 m)
1992–2005 0.11 [0.01 to 0.21]
b
Deep hydrographic sections (Purkey and Johnson, 2010) Trend only
Thermosteric Component
(whole depth)
1971–2010 0.8 [0.5 to 1.1]
c
Combination of estimates from 0 to 700 m, 700 to 2000 m, and below 2000 m
c
Trend only
1993–2010 1.1 [0.8 to 1.4]
c
Combination of estimates from 0–700 m, 700 to 2000 m, and below 2000 m
c
Trend only
Table 3.1 | Estimated trends in GMSL and components over different periods from representative time-series. Trends and uncertainty have been estimated from a time series
provided by the authors using ordinary least squares with the uncertainty representing the 90% confidence interval. The model fit for yearly averaged time series was a bias + trend;
the model fit for monthly and 10-day averaged data was a bias + trend + seasonal sinusoids. Uncertainty accounts for correlations in the residuals.
Notes:
a
Uncertainty estimated from fit to Nerem et al. (2010) time series and includes potential systematic error owing to drift of altimeter, estimated to be ±0.4 mm yr
–1
(Beckley et al., 2010; Nerem
et al., 2010), applied as the root-sum-square (RSS) with the least squares error estimate. The uncertainty in drift contains uncertainty in the reference frame, orbit and instrument.
b
Trend value taken from Purkey and Johnson (2010), Table 1. Uncertainty represents the 2.5–97.5% confidence interval.
c
Assumes no trend below 2000 m before 1 January 1992, then value from Purkey and Johnson (2010) afterwards. Uncertainty for 0 to 700 m, 700 to 2000 m and below 2000 m is assumed to
be uncorrelated, and uncertainty is calculated as RSS of the uncertainty for each layer.
292
Chapter 3 Observations: Ocean
3
also provides an important sink for carbon dioxide released by human
activities, the anthropogenic CO
2
(C
ant
). Currently, an amount of CO
2
equivalent to approximately 30% of the total human emissions of CO
2
to the atmosphere is accumulating in the ocean (Mikaloff-Fletcher et
al., 2006; Le Quéré et al., 2010). In this section, observations of change
in the ocean uptake of carbon, the inventory of C
ant
, and ocean acidifi-
cation are assessed, as well as changes in oxygen and nutrients. Chap-
ter 6 provides a synthesis of the overall carbon cycle, including the
ocean, atmosphere and biosphere and considering both past trends
and future projections.
3.8.1 Carbon
3.8.1.1 Ocean Uptake of Carbon
The air–sea flux of CO
2
is computed from the observed difference in the
partial pressure of CO
2
(pCO
2
) across the air–water interface (∆pCO
2
= pCO
2,sw
- pCO
2,air
), the solubility of CO
2
in seawater, and the gas
transfer velocity (Wanninkhof et al., 2009). However, the limited geo-
graphic and temporal coverage of the ∆pCO
2
measurement as well as
uncertainties in wind forcing and transfer velocity parameterizations
mean that uncertainties in global and regional fluxes calculated from
measurements of ∆pCO
2
can be as larges as ±50% (Wanninkhof et al.,
2013). Using ∆pCO
2
data in combination with the riverine input Gruber
et al. (2009) estimated a global uptake rate of 1.9 [1.2 to 2.5] PgC
yr
–1
for the time period 1995–2000 and Takahashi et al. (2009) found
2.0 [1.0 to 3.0] PgC yr
–1
normalized to the year 2000. Uncertainties in
fluxes calculated from ∆pCO
2
are too large to detect trends in global
ocean carbon uptake.
Trends in surface ocean pCO
2
are calculated from ocean time series sta-
tions and repeat hydrographic sections in the North Atlantic and North
Pacific (Table 3.2). At all locations and for all time periods shown, pCO
2
in both the atmosphere and ocean has increased, while pH and [CO
3
2–
]
have decreased. At some sites, oceanic surface pCO
2
increased faster
than the atmospheric trend, implying a decreasing uptake of atmos-
pheric CO
2
at those locations. The oceanic pCO
2
trend can differ from
that in the atmosphere owing to changes in the intensity of biological
production and changes in physical conditions, for instance between
El Niño and La Niña (Keeling et al., 2004; Midorikawa et al., 2005;
Yoshikawa-Inoue and Ishii, 2005; Takahashi et al., 2006, 2009; Schuster
and Watson, 2007; Ishii et al., 2009; McKinley et al., 2011; Bates, 2012;
Lenton et al., 2012).
Although local variations of ∆pCO
2
with time have little effect on the
atmospheric CO
2
growth rate in the short term, they provide impor-
tant information on the dynamics of the ocean carbon cycle and the
potential for longer-term climate feedbacks. For example, El Niño and
La Niña can drive large changes in the efflux of CO
2
in the Pacific.
Differences in ∆pCO
2
can exceed 100 µatm in the eastern and central
equatorial Pacific between El Niño and La Niña; an increase in ∆pCO
2
observed between 1998 and 2004 was attributed to wind and circula-
tion changes associated with the Pacific Decadal Oscillation (Feely et
al., 2006). CO
2
uptake in the North Atlantic decreased by 0.24 [0.19–
0.29] PgC yr
–1
between 1994 and 2003 (Schuster and Watson, 2007)
and has partially recovered since then (Watson et al., 2009). Linear
trends for the North Atlantic from 1995 to 2009 reveal an increased
uptake (Schuster et al., 2013). Uptake of CO
2
in the Subtropical Mode
Water (STMW) of the North Atlantic was enhanced during the 1990s,
a predominantly positive phase of the NAO, and much reduced in the
2000s when the NAO phase was neutral or negative (Bates, 2012).
Observations in the Indian and Pacific sectors of the Southern Ocean
were interpreted as evidence for reduced winter-time CO
2
uptake as a
result of increased winds, increased upwelling and outgassing of natu-
ral CO
2
(Metzl, 2009; Lenton et al., 2012).
3.8.1.2 Changes in the Oceanic Inventory of Anthropogenic
Carbon Dioxide
Ocean carbon uptake and storage is inferred from changes in the
inventory of anthropogenic carbon. C
ant
cannot be measured direct-
ly but is calculated from observations of ocean properties (Appendix
3.A discusses the sampling on which the ocean carbon inventory is
based). Two independent data-based methods to calculate anthropo-
genic carbon inventories exist: the ΔC* method (Sabine et al., 2004),
and the transit time distribution (TTD) method (Waugh et al., 2006).
The Green’s function approach that applies the maximum entropy
de-convolution methodology (Khatiwala et al., 2009) is related to the
latter. These approaches use different tracer data, for instance, the TTD
method is based mostly on chlorofluorcarbon measurements. Changes
due to variability in ocean productivity (Chavez et al., 2011) are not
considered.
Estimates of the global inventory of C
ant
(including marginal seas) cal-
culated using these methods have a mean value of 118 PgC and a
range of 93 to 137 PgC in 1994 and a mean of 160 PgC and range of
134 to 186 PgC in 2010 (Sabine et al., 2004; Waugh et al., 2006; Khati-
wala et al., 2009, 2013). When combined with model results (Mikaloff-
Fletcher et al., 2006; Doney et al., 2009; Gerber et al., 2009; Graven
et al., 2012), Khatiwala et al. (2013) arrive at a “best” estimate of
the global ocean inventory (including marginal seas) of anthropogenic
carbon from 1750 to 2010 of 155 PgC with an uncertainty of ±20%
(Figure 3.16). While the estimates of total inventory agree within their
uncertainty, the different methods result in significant differences in
the inferred spatial distribution of C
ant
, particularly at high latitudes.
The C
ant
inventory “best” estimate of 155 PgC (Khatiwala et al., 2013;
Figure 3.16) corresponds to an uptake rate of 2.3 (range of 1.7 to 2.9)
PgC yr
–1
from 2000 to 2010, in close agreement with an independent
estimate of 2.5 (range of 1.8 to 3.2) PgC yr
–1
based on atmospheric O
2
/
N
2
measurements obtained for the same period (Ishidoya et al., 2012).
The O
2
/N
2
method resulted in 2.2 ± 0.6 PgC yr
–1
for the time period
1990 to 2000 and 2.5 ± 0.6 for the period from 2000 to 2010 (Keeling
and Manning, 2014). These estimates are also consistent with an inde-
pendent estimate of 1.9 ± 0.4 PgC yr
–1
for the period between 1970
and 1990 based on depth-integrated d
13
C changes (Quay et al., 2003)
and with estimates inferred from ∆pCO
2
.
The storage rate of anthropogenic CO
2
is assessed by calculating the
change in C
ant
concentrations between two time periods. Regional
observations of the storage rate are in general agreement with that
expected from the increase in atmospheric CO
2
concentrations and
with the tracer-based estimates. However, there are significant spatial
and temporal variations in the degree to which the inventory of C
ant
293
3
Observations: Ocean Chapter 3
tracks changes in the atmosphere (Figure 3.17). The North Atlantic, in
particular, is an area with high variability in circulation and deep water
formation, influencing the C
ant
inventory. As a result of the decline in
Labrador Sea Water (LSW) formation since 1997 (Rhein et al., 2011),
the C
ant
increase between 1997 and 2003 was smaller in the subpolar
North Atlantic than expected from the atmospheric increase, in con-
trast to the subtropical and equatorial Atlantic (Steinfeldt et al., 2009).
Perez et al. (2010) also noted the dependence of the C
ant
storage rate
in the North Atlantic on the NAO, with high C
ant
storage rate during
phases of high NAO (i.e., high LSW formation rates) and low storage
during phases of low NAO (low formation). Wanninkhof et al. (2010)
found a smaller inventory increase in the North Atlantic compared to
the South Atlantic between 1989 and 2005.
Ocean observations are insufficient to assess whether there has been
a change in the rate of total (anthropogenic plus natural) carbon
uptake by the global ocean. Evidence from regional ocean studies
(often covering relatively short time periods), atmospheric observa-
tions and models is equivocal, with some studies suggesting the ocean
uptake rate of total CO
2
may have declined (Le Quéré et al., 2007;
Schuster and Watson, 2007; McKinley et al., 2011) while others con-
clude that there is little evidence for a decline (Knorr, 2009; Gloor et
al., 2010; Sarmiento et al., 2010). A study based on atmospheric CO
2
observations and emission inventories concluded that global carbon
uptake by land and oceans doubled from 1960 to 2010, implying
that it is unlikely that on a global scale both land and ocean sinks
decreased (Ballantyne et al., 2012).
In summary, the high agreement between multiple lines of independ-
ent evidence for increases in the ocean inventory of C
ant
underpins the
conclusion that it is virtually certain that the ocean is sequestering
anthropogenic carbon dioxide and very likely that the oceanic C
ant
inventory increased from 1994 to 2010. Oceanic carbon uptake rates
calculated using different data sets and methods agree within their
uncertainties and very likely range between 1.0 and 3.2 PgC yr
–1
.
3.8.2 Anthropogenic Ocean Acidification
The uptake of CO
2
by the ocean changes the chemical balance of
seawater through the thermodynamic equilibrium of CO
2
with sea-
water. Dissolved CO
2
forms a weak acid (H
2
CO
3
) and, as CO
2
in sea-
water increases, the pH, carbonate ion (CO
3
2–
), and calcium carbonate
(CaCO
3
) saturation state of seawater decrease while bicarbonate ion
(HCO
3
–
) increases (FAQ 3.3). Variations in oceanic total dissolved inor-
ganic carbon (C
T
= CO
2
+ CO
3
2–
+ HCO
3
–
) and pCO
2
reflect changes in
both the natural carbon cycle and the uptake of anthropogenic CO
2
from the atmosphere. The mean pH (total scale) of surface waters
ranges between 7.8 and 8.4 in the open ocean, so the ocean remains
mildly basic (pH > 7) at present (Orr et al., 2005a; Feely et al., 2009).
Ocean uptake of CO
2
results in gradual acidification of seawater; this
0
0.4
0.8
1.2
1.6
2
2.4
60°S
30°S
0°
30°N
60°N
(mol m
-2
y
-1
)(mol m
-2
y
-1
)(mol m
-2
y
-1
)
60°S
30°S
0°
30°N
60°N
60°S
30°S
0°
30°N
60°N
0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
0.8
0
0.1
0.2
0.3
0.4
0.5
0.6
0.7
0.8
0.9
Atlantic Ocean Pacific OceanIndian Ocean
75°W
50°W
25°W
0°
25°E
120°
E
150°E
150°W
120°W
90°W
180°W
25°E
50°E
100°E
125°E
150°
E
75°E
Figure 3.17 | Maps of storage rate distribution of anthropogenic carbon (mol m
–2
yr
–1
) for the three ocean basins (left to right: Atlantic, Pacific and Indian Ocean) averaged over
1980–2005 estimated by the Green’s function approach (Khatiwala et al., 2009). Note that a different colour scale is used in each basin.
60
o
E 120
o
E 180
o
W 120
o
W 60
o
W 0
o
80
o
S
40
o
S
0
o
40
o
N
80
o
N
0
20
40
60
80
100
120
140
160
150
o
W
120
o
W
90
o
W
60
o
W
30
o
W
0
o
30
o
E
60
o
E
90
o
E
120
o
E
150
o
E
180
o
W
o
80
o
70
N
N
o
(mol m
-2
)
Figure 3.16 | Compilation of the 2010 column inventories (mol m
–2
) of anthropogenic CO
2
: the global Ocean excluding the marginal seas (updated from Khatiwala et al., 2009)
150 ± 26 PgC; Arctic Ocean (Tanhua et al., 2009) 2.7 to 3.5 PgC; the Nordic Seas (Olsen et al., 2010) 1.0 to 1.6 PgC; the Mediterranean Sea (Schneider et al., 2010) 1.6 to 2.5
PgC; the Sea of Japan(Park et al., 2006) 0.40 ± 0.06 PgC. From Khatiwala et al. (2013).
294
Chapter 3 Observations: Ocean
3
process is termed ocean acidification (Box 3.2) (Broecker and Clark,
2001; Caldeira and Wickett, 2003). The observed decrease in ocean
pH of 0.1 since the beginning of the industrial era corresponds to a
26% increase in the hydrogen ion concentration [H
+
]
concentration of
seawater (Orr et al., 2005b; Feely et al., 2009). The consequences of
changes in pH, CO
3
2–
, and the saturation state of CaCO
3
minerals for
marine organisms and ecosystems are just beginning to be understood
(see WGII Chapters 5, 6, 28 and 30).
A global mean decrease in surface water pH of 0.08 from 1765 to 1994
was calculated based on the inventory of anthropogenic CO
2
(Sabine
et al., 2004), with the largest reduction (–0.10) in the northern North
Atlantic and the smallest reduction (–0.05) in the subtropical South
Pacific. These regional variations in the size of the pH decrease are con-
sistent with the generally lower buffer capacities of the high latitude
oceans compared to lower latitudes (Egleston et al., 2010).
Direct measurements on ocean time-series stations in the North Atlan-
tic and North Pacific record decreasing pH with rates ranging between
–0.0014 and –0.0024 yr
–1
(Table 3.2, Figure 3.18; Bates, 2007, 2012;
Santana-Casiano et al., 2007; Dore et al., 2009; Olafsson et al., 2009;
González-Dávila et al., 2010). Directly measured pH differences in the
surface mixed layer along repeat transects in the central North Pacific
Ocean between Hawaii and Alaska showed a –0.0017 yr
–1
decline in
pH between 1991 and 2006, in agreement with observations at the
time-series sites (Byrne et al., 2010). This rate of pH change is also
consistent with repeat transects of CO
2
and pH measurements in the
western North Pacific (winter: –0.0018 ± 0.0002 yr
–1
; summer: –0.0013
± 0.0005 yr
–1
) (Midorikawa et al., 2010). The pH changes in southern
ocean surface waters are less certain because of the paucity of long-
term time-series observations there, but pCO
2
measurements collect-
ed by ships-of-opportunity indicate similar rates of pH decrease there
(Takahashi et al., 2009).
Uptake of anthropogenic CO
2
is the dominant cause of observed
changes in the carbonate chemistry of surface waters (Doney et al.,
2009). Changes in carbonate chemistry in subsurface waters can also
reflect local physical and biological variability. As an example, while
pH changes in the mixed layer of the North Pacific Ocean can be
explained solely by equilibration with atmospheric CO
2
, declines in pH
between 800 m and the mixed layer in the time period 1991–2006
were attributed in approximately equal measure to anthropogenic and
natural variations (Byrne et al., 2010). Figure 3.19 shows the portion
of pH changes between the surface and 1000 m that were attributed
solely to the effects of anthropogenic CO
2
. Seawater pH and [CO
3
2–
]
decreased by 0.0014 to 0.0024 yr
–1
and ~0.4 to 0.9 µmol kg
–1
yr
–1
,
respectively, between 1988 and 2009 (Table 3.2). Over longer time
periods, anthropogenic changes in ocean chemistry are expected to
become increasingly prominent relative to changes imparted by physi-
cal and biological variability.
The consistency of these observations demonstrates that the pH of
surface waters has decreased as a result of ocean uptake of anthropo-
genic CO
2
from the atmosphere. There is high confidence that the pH
decreased by 0.1 since the preindustrial era.
270
300
330
360
390
420
1985 1990 1995 2000 2005 2010
pCO
2
(µatm)
8.05
8.10
8.15
8.20
1985 1990 1995 2000 2005 2010
in situ pH (total scale)
210
230
250
270
1985 1990 1995 2000 2005 2010
Y
ea
r
[CO
3
2−
] (μmol kg
−1
)
BATS
ALOHA
ESTOC
BATS
ALOHA
ESTOC
BATS
ALOHA
ESTOC
Figure 3.18 | Long-term trends of surface seawater pCO
2
(top), pH (middle) and car-
bonate ion (bottom) concentration at three subtropical ocean time series in the North
Atlantic and North Pacific Oceans, including (a) Bermuda Atlantic Time-series Study
(BATS, 31°40’N, 64°10’W; green) and Hydrostation S (32°10’, 64°30’W) from 1983
to present (updated from Bates, 2007); (b) Hawaii Ocean Time-series (HOT) at Station
ALOHA (A Long-term Oligotrophic Habitat Assessment; 22°45’N, 158°00’W; orange)
from 1988 to present (updated from Dore et al., 2009) and (c) European Station for
Time series in the Ocean (ESTOC, 29°10’N, 15°30’W; blue) from 1994 to present
(updated from González-Dávila et al., 2010). Atmospheric pCO
2
(black) from the Mauna
Loa Observatory Hawaii is shown in the top panel. Lines show linear fits to the data,
whereas Table 3.2 give results for harmonic fits to the data (updated from Orr, 2011).
3.8.3 Oxygen
As a consequence of the early introduction of standardized methods
and the relatively wide interest in the distribution of dissolved oxygen,
the historical record of marine oxygen observations is generally richer
than that of other biogeochemical parameters, although still sparse
compared to measurements of temperature and salinity (Appendix
3.A). Dissolved oxygen changes in the ocean thermocline has generally
decreased since 1960, but with strong regional variations (Keeling et
al., 2010; Keeling and Manning, 2014). Oxygen concentrations at 300
dbar decreased between 50°S and 50°N at a mean rate of 0.63 µmol
kg
–1
per decade between 1960 and 2010 (Stramma et al., 2012). For
the period 1970 to 1990, the mean annual global oxygen loss between
100 m and 1000 m was calculated to be 0.55 ± 0.13 × 10
14
mol yr
–1
(Helm et al., 2011).
295
3
Observations: Ocean Chapter 3
Box 3.2 | Ocean Acidification
Ocean acidification refers to a reduction in pH of the ocean over an extended period, typically decades or longer, caused primarily by the
uptake of carbon dioxide (CO
2
) from the atmosphere. Ocean acidification can also be caused by other chemical additions or subtractions
from the oceans that are natural (e.g., increased volcanic activity, methane hydrate releases, long-term changes in net respiration) or
human-induced (e.g., release of nitrogen and sulphur compounds into the atmosphere). Anthropogenic ocean acidification refers to the
component of pH reduction that is caused by human activity (IPCC, 2011).
Since the beginning of the industrial era, the release of CO
2
from
industrial and agricultural activities has resulted in atmospheric
CO
2
concentrations that have increased from approximately 280
ppm to about 392 ppm in 2012 (Chapter 6). The oceans have
absorbed approximately 155 PgC from the atmosphere over
the last two and a half centuries (Sabine et al., 2004; Khatiwala
et al., 2013). This natural process of absorption has benefited
humankind by significantly reducing the greenhouse gas levels
in the atmosphere and abating some of the impacts of global
warming. However, the ocean’s uptake of carbon dioxide is
having a significant impact on the chemistry of seawater. The
average pH of ocean surface waters has already fallen by about
0.1 units, from about 8.2 to 8.1 (total scale), since the beginning
of the industrial revolution (Orr et al., 2005a; Figure 1; Feely et
al., 2009). Estimates of future atmospheric and oceanic carbon
dioxide concentrations indicate that, by the end of this century,
the average surface ocean pH could be lower than it has been
for more than 50 million years (Caldeira and Wickett, 2003).
The major controls on seawater pH are atmospheric CO
2
exchange, the production and respiration of dissolved and
particulate organic matter in the water column, and the formation
and dissolution of calcium carbonate minerals. Oxidation of
organic matter lowers dissolved oxygen concentrations, adds
CO
2
to solution, reduces pH, carbonate ion (CO
3
2–
) and calcium
carbonate (CaCO
3
) saturation states (Box 3.2, Figure 2), and
lowers the pH of seawater in subsurface waters (Byrne et al.,
2010). As a result of these processes, minimum pH values in the
oceanic water column are generally found near the depths of the
oxygen minimum layer. When CO
2
reacts with seawater it forms
carbonic acid (H
2
CO
3
), which is highly reactive and reduces the
concentration of carbonate ion (Box 3.2, Figure 2) and can affect
shell formation for marine animals such as corals, plankton,
and shellfish. This process could affect fundamental biological
and chemical processes of the sea in coming decades (Fabry et
al., 2008; Doney et al., 2009; WGII Chapters 5, 6, 28 and 30).
(continued on next page)
Box 3.2, Figure 1 | National Center for Atmospheric Research Community
Climate System Model 3.1 (CCSM3)-modeled decadal mean pH at the sea surface
centred on the years 1875 (top) and 1995 (middle). Global Ocean Data Analysis
Project (GLODAP)-based pH at the sea surface, nominally for 1995 (bottom).
Deep and shallow-water coral reefs are indicated with magenta dots. White areas
indicate regions with no data. (After Feely et al., 2009.)
1875 Model
1995 Model
GLODAP Observations
8.40
8.35
8.30
8.25
8.20
8.15
8.05
8.10
8.00
7.95
7.90
pH
The long-term deoxygenation of the open ocean thermocline is consist-
ent with the expectation that warmer waters can hold less dissolved
oxygen (solubility effect), and that warming-induced stratification
leads to a decrease in the transport of dissolved oxygen from surface
to subsurface waters (stratification effect) (Matear and Hirst, 2003;
Deutsch et al., 2005; Frölicher et al., 2009). Observations of oxygen
change suggested that about 15% of the oxygen decline between
1970 and 1990 could be explained by warming and the remainder by
reduced ventilation due to increased stratification (Helm et al., 2011;
see Table 6.14).
Oxygen concentrations in the tropical ocean thermocline decreased
in each of the ocean basins over the last 50 years (Ono et al., 2001;
Stramma et al., 2008; Keeling et al., 2010; Helm et al., 2011), resulting
in an expansion of the dissolved oxygen minimum zones. A comparison
of data between 1960 and 1974 with those from 1990 to 2008 showed
296
Chapter 3 Observations: Ocean
3
that oxygen concentrations decreased in most tropical regions at an
average rate of 2 to 3 µmol kg
–1
per decade (Figure 3.20; Stramma et
al., 2010). Data from one of the longest time-series sites in the subpo-
lar North Pacific (Station Papa, 50°N, 145°W) reveal a persistent declin-
ing oxygen trend in the thermocline over the last 50 years (Whitney et
al., 2007), superimposed on oscillations with time scales of a few years
to two decades. Stendardo and Gruber (2012) found dissolved oxygen
decreases in upper water masses of the North Atlantic and increases
in intermediate water masses. The changes were caused by changes in
solubility as well as changes in ventilation and circulation over time.
In contrast to the widely distributed oxygen declines, oxygen increased
in the thermoclines of the Indian and South Pacific Oceans from the
1990s to the 2000s (McDonagh et al., 2005; Álvarez et al., 2011),
apparently due to strengthened circulation driven by stronger winds
(Cai, 2006; Roemmich et al., 2007). In the southern Indian Ocean below
the thermocline, east of 75°E, oxygen decreased between 1960 and
2010 most prominently on the isopycnals σ
q
= 26.9 to 27.0 (Kobayas-
hi et al., 2012). While some studies suggest a widespread decline of
oxygen in the Southern Ocean (e.g., Helm et al., 2011), other studies
Box 3.2, Figure 2 | Distribution of (a) pH and (b) carbonate (CO
3
2–
) ion concentration in the Pacific, Atlantic and Indian Oceans. The data are from the World Ocean
Circulation Experiment/Joint Global Ocean Flux Study/Ocean Atmosphere Carbon Exchange Study global carbon dioxide (CO
2
) survey (Sabine et al., 2005). The lines
show the mean pH (red solid line, top panel), mean CO
3
2–
(red solid line, bottom panel), and aragonite and calcite (black solid and dashed lines, bottom panel) satura-
tion values for each of these basins (modified from Feely et al., 2009). The shaded areas show the range of values within the ocean basins. Dissolution of aragonite
and calcite shells and skeletons occurs when CO
3
2–
concentrations drop below the saturation level, reducing the ability of calcifying organisms to produce their shells
and skeletons.
Pacific
Atlantic
Indian
(a)
0
200
400
600
800
1000
Depth (m)
2000
4000
6000
(b)
0
200
400
600
800
1000
2000
4000
6000
Depth (m)
7.6
7.8
8.2
7.6
7.8
8.0
8.2
7.6
7.8
8.0
8.2
50
150
250
50
150 250 50
150
250
8.0
pH
(total scale)
CO
3
2-
(µmol kg
-1
)
Calcite
saturation
Aragonite
saturation
pH (total scale)
Carbonate ion concentration (µmol kg
-1
)
Box 3.2 (continued)
Figure 3.19 | ΔpH
ant
: pH change attributed to the uptake of anthropogenic carbon
between 1991 and 2006, at about 150°W, Pacific Ocean (from Byrne et al., 2010). The
red lines show the layers of constant density.
CLIVAR P16N (2006) - WOCE P16N (1991): ∆pH (anthropogenic)
250
750
500
1000
Depth (m)
25°
30°
35°
40°
45°
50°
55°N
0
27.4
26.8
27.2
27
26.6
26.4
26.2
26
25.8
25.6
25.4
25.2
24.8
0.01
0
0
-0.01
-0.01
-0.02
-0.03
0
0.08
0.06
0.04
0.02
0
-0.02
-0.04
-0.06
-0.08
Alaska
Hawaii
pH
0
show regions of alternating sign (e.g., Stramma et al., 2010), reflecting
differences in data and period considered.
297
3
Observations: Ocean Chapter 3
Frequently Asked Questions
FAQ 3.3 | How Does Anthropogenic Ocean Acidification Relate to Climate Change?
Both anthropogenic climate change and anthropogenic ocean acidification are caused by increasing carbon dioxide
concentrations in the atmosphere. Rising levels of carbon dioxide (CO
2
), along with other greenhouse gases, indi-
rectly alter the climate system by trapping heat as it is reflected back from the Earth’s surface. Anthropogenic ocean
acidification is a direct consequence of rising CO
2
concentrations as seawater currently absorbs about 30% of the
anthropogenic CO
2
from the atmosphere.
Ocean acidification refers to a reduction in pH over an extended period, typically decades or longer, caused primari-
ly by the uptake of CO
2
from the atmosphere. pH is a dimensionless measure of acidity. Ocean acidification describes
the direction of pH change rather than the end point; that is, ocean pH is decreasing but is not expected to become
acidic (pH < 7). Ocean acidification can also be caused by other chemical additions or subtractions from the oceans
that are natural (e.g., increased volcanic activity, methane hydrate releases, long-term changes in net respiration)
or human-induced (e.g., release of nitrogen and sulphur compounds into the atmosphere). Anthropogenic ocean
acidification refers to the component of pH reduction that is caused by human activity.
Since about 1750, the release of CO
2
from industrial and agricultural activities has resulted in global average atmo-
spheric CO
2
concentrations that have increased from 278 to 390.5 ppm in 2011. The atmospheric concentration of
CO
2
is now higher than experienced on the Earth for at least the last 800,000 years and is expected to continue to
rise because of our dependence on fossil fuels for energy. To date, the oceans have absorbed approximately 155 ±
30 PgC from the atmosphere, which corresponds to roughly one-fourth of the total amount of CO
2
emitted (555 ±
85 PgC) by human activities since preindustrial times. This natural process of absorption has significantly reduced
the greenhouse gas levels in the atmosphere and minimized some of the impacts of global warming. However, the
ocean’s uptake of CO
2
is having a significant impact on the chemistry of seawater. The average pH of ocean surface
waters has already fallen by about 0.1 units, from about 8.2 to 8.1 since the beginning of the Industrial Revolution.
Estimates of projected future atmospheric and oceanic CO
2
concentrations indicate that, by the end of this century,
the average surface ocean pH could be 0.2 to 0.4 lower than it is today. The pH scale is logarithmic, so a change of
1 unit corresponds to a 10-fold change in hydrogen ion concentration.
When atmospheric CO
2
exchanges across the air–sea interface it reacts with seawater through a series of four chem-
ical reactions that increase the concentrations of the carbon species: dissolved carbon dioxide (CO
2(aq)
), carbonic acid
(H
2
CO
3
) and bicarbonate (HCO
3
–
):
CO
2(atmos)
CO
2(aq)
(1)
CO
2(aq)
+ H
2
O H
2
CO
3
(2)
H
2
CO
3
H
+
+ HCO
3
–
(3)
HCO
3
–
H
+
+ CO
3
2–
(4)
Hydrogen ions (H
+
) are produced by these reactions. This increase in the ocean’s hydrogen ion concentration cor-
responds to a reduction in pH, or an increase in acidity. Under normal seawater conditions, more than 99.99% of
the hydrogen ions that are produced will combine with carbonate ion (CO
3
2–
) to produce additional HCO
3
–
. Thus,
the addition of anthropogenic CO
2
into the oceans lowers the pH and consumes carbonate ion. These reactions are
fully reversible and the basic thermodynamics of these reactions in seawater are well known, such that at a pH of
approximately 8.1 approximately 90% the carbon is in the form of bicarbonate ion, 9% in the form of carbonate
ion, and only about 1% of the carbon is in the form of dissolved CO
2
. Results from laboratory, field, and modeling
studies, as well as evidence from the geological record, clearly indicate that marine ecosystems are highly suscep-
tible to the increases in oceanic CO
2
and the corresponding decreases in pH and carbonate ion.
Climate change and anthropogenic ocean acidification do not act independently. Although the CO
2
that is taken up
by the ocean does not contribute to greenhouse warming, ocean warming reduces the solubility of carbon dioxide
in seawater; and thus reduces the amount of CO
2
the oceans can absorb from the atmosphere. For example, under
doubled preindustrial CO
2
concentrations and a 2°C temperature increase, seawater absorbs about 10% less CO
2
(10% less total carbon, C
T
) than it would with no temperature increase (compare columns 4 and 6 in Table 1), but
the pH remains almost unchanged. Thus, a warmer ocean has less capacity to remove CO
2
from the atmosphere, yet
still experiences ocean acidification. The reason for this is that bicarbonate is converted to carbonate in a warmer
ocean, releasing a hydrogen ion thus stabilizing the pH. (continued on next page)
298
Chapter 3 Observations: Ocean
3
FAQ 3.3, Figure 1 | A smoothed time series of atmospheric CO
2
mole fraction (in ppm) at the atmospheric Mauna Loa Observatory (top red line), surface ocean
partial pressure of CO
2
(pCO
2
; middle blue line) and surface ocean pH (bottom green line) at Station ALOHA in the subtropical North Pacific north of Hawaii for the
period from1990–2011 (after Doney et al., 2009; data from Dore et al., 2009).The results indicate that the surface ocean pCO
2
trend is generally consistent with the
atmospheric increase but is more variable due to large-scale interannual variability of oceanic processes.
Parameter
Pre-industrial
(280 ppmv)
20°C
2 × Pre-industrial
(560 ppmv)
20°C
(% change relative
to pre-industrial)
2 × Pre-industrial
(560 ppmv)
22°C
(% change relative
to pre-industrial)
pH 8.1714 7.9202 – 7.9207 –
H
+
(mol kg
–1
) 6.739e
–9
1.202e
–8
(78.4) 1.200e
–8
(78.1)
CO
2(aq)
(µmol kg
–1
) 9.10 18.10 (98.9) 17.2 (89.0)
HCO
3
–
(µmol kg
–1
) 1723.4 1932.8 (12.15) 1910.4 (10.9)
CO
3
2–
(µmol kg
–1
) 228.3 143.6 (-37.1) 152.9 (–33.0)
C
T
(µmol kg
–1
) 1960.8 2094.5 (6.82) 2080.5 (6.10)
Notes:
a
CO
2(aq)
= dissolved CO
2
, H
2
CO
3
= carbonic acid, HCO
3
–
= bicarbonate, CO
3
2–
= carbonate, C
T
= total carbon = CO
2(aq)
+ HCO
3
–
+ CO
3
2–
).
FAQ 3.3, Table 1 | Oceanic pH and carbon system parameter changes in surface water for a CO
2
doubling from the preindustrial atmosphere without and with a
2°C warming
a
.
8.30
8.25
8.20
8.15
8.10
8.05
8.00
400
375
350
325
300
275
1990 1992 1994 19961998 2000 2002 2004 2006 2008 2010 2012
CO
2
Time Series in the North Pacific
Year
pH
pCO
2
(µatm) CO
2
(ppm)
160°W 158°W 156°W
23°N
22°N
21°N
20°N
19°N
Station Mauna Loa
Station Aloha
FAQ 3.3 (continued)
Coastal regions have also experienced long-term dissolved oxygen
changes. Bograd et al. (2008) reported a substantial reduction of the
thermocline oxygen content in the southern part of the California Cur-
rent from 1984 to 2002, resulting in a shoaling of the hypoxic bound-
ary (marked by oxygen concentrations of about 60 µmol kg
–1
). Off the
British Columbia coast, oxygen concentrations in the near bottom
waters decreased an average of 1.1 µmol kg
–1
yr
–1
over a 30-year
period (Chan et al., 2008). These changes along the west coast of North
America appear to have been largely caused by the open ocean dis-
solved oxygen decrease and local processes associated with decreased
vertical dissolved oxygen transport following near-surface warming
and increased stratification. Gilbert et al. (2010) found evidence that
for the time period 1976–2000 oxygen concentrations between 0 and
300 m depth were declining about 10 times faster in the coastal ocean
than in the open ocean, and an increase in the number of hypoxic
zones was observed since the 1960s (Diaz and Rosenberg, 2008).
3.8.4 Nutrients
Nutrient concentrations in the surface ocean surface are influenced by
human impacts on coastal runoff and on atmospheric deposition, and
by changing nutrient supply from the ocean’s interior into the mixed
layer (for instance due to increased stratification). Changing nutrient
distributions might influence the magnitude and variability of the
ocean’s biological carbon pump.
Globally, the manufacture of nitrogen fertilizers has continued to
increase (Galloway et al., 2008) accompanied by increasing eutrophi-
299
3
Observations: Ocean Chapter 3
Site Period
pCO
2
atm
(μatm yr
–1
)
pCO
2
sea
(μatm yr
–1
)
pH
*
(yr
–1
)
[CO
3
2–
]
(μmol kg
–1
yr
–1
)
Ω
a
(yr
–1
)
a. Published trends
BATS
1983–2005
a
1.78 ± 0.02 1.67 ± 0.28
–0.0017 ± 0.0003 –0.47 ± 0.09
–0.007 ± 0.002
1983–2005
b
1.80 ± 0.02 1.80 ± 0.13 –0.0017 ± 0.0001 –0.52 ± 0.02 –0.006 ± 0.001
ALOHA
1988–2007
c
1.68 ± 0.03 1.88 ± 0.16
–0.0019 ± 0.0002 —
–0.0076 ± 0.0015
1998–2007
d
—
—
–0.0014 ± 0.0002 —
—
ESTOC
1995–2004
e
— 1.55 ± 0.43
–0.0017 ± 0.0004 —
—
1995–2004
f
1.6 ± 0.7 1.55
–0.0015 ± 0.0007 –0.90 ± 0.08
–0.0140 ± 0.0018
IS
1985–2006
g
1.69 ± 0.04 2.15 ± 0.16
–0.0024 ± 0.0002 —
–0.0072 ± 0.0007
g
N. Pacific
1991–2006
h
— —
–0.0017 —
—
N. Pacific
1983–2008
i
Summer 1.54 ± 0.08
Winter 1.65 ± 0.05
Summer 1.37 ± 0.33
Winter 1.58 ± 0.12
Summer –0.0013 ± 0.0005
Winter –0.0018 ± 0.0002
—
—
Coast of western
N. Pacific
1994–2008
k
1.99 ± 0.02 1.54 ± 0.33 –0.0020 ± 0.0007 — –0.012 ± 0.005
b. Updated trends
j,l
BATS
1983–2009 1.66 ± 0.01 1.92 ± 0.08
–0.0019 ± 0.0001 –0.59 ± 0.04
–0.0091 ± 0.0006
1985–2009 1.67 ± 0.01 2.02 ± 0.08
–0.0020 ± 0.0001 –0.68 ± 0.04
–0.0105 ± 0.0006
1988–2009 1.73 ± 0.01 2.22 ± 0.11
–0.0022 ± 0.0001 –0.87 ± 0.05
–0.0135 ± 0.0008
1995–2009 1.90 ± 0.01 2.16 ± 0.18
–0.0021 ± 0.0002 –0.80 ± 0.08
–0.0125 ± 0.0013
ALOHA
1988–2009 1.73 ± 0.01 1.82 ± 0.07
–0.0018 ± 0.0001 –0.52 ± 0.04
–0.0083 ± 0.0007
1995–2009 1.92 ± 0.01 1.58 ± 0.13
–0.0015 ± 0.0001 –0.40 ± 0.07
–0.0061 ± 0.0028
ESTOC
1995–2009 1.88 ± 0.02 1.83 ± 0.15
–0.0017 ± 0.0001 –0.72 ± 0.05
–0.0123 ± 0.0015
IS
1985–2009
m
1.75 ± 0.01 2.07 ± 0.15
–0.0024 ± 0.0002 –0.47 ± 0.04
–0.0071 ± 0.0006
1988–2009
m
1.70 ± 0.01 1.96 ± 0.22
–0.0023 ± 0.0003 –0.48 ± 0.05
–0.0073 ± 0.0008
1995–2009
m
1.90 ± 0.01 2.01 ± 0.37
–0.0022 ± 0.0004 –0.40 ± 0.08
–0.0062 ± 0.0012
Table 3.2 | Published and updated long-term trends of atmospheric (pCO
2
atm
) and seawater carbonate chemistry (i.e., surface-water pCO
2
, and corresponding calculated pH, CO
3
2–
,
and aragonite saturation state (Ωa) at four ocean time series in the North Atlantic and North Pacific oceans: (1) Bermuda Atlantic Time-series Study (BATS, 31°40’N, 64°10’W)
and Hydrostation S (32°10’N, 64°30’W) from 1983 to present (Bates, 2007); (2) Hawaii Ocean Time series (HOT) at Station ALOHA (A Long-term Oligotrophic Habitat Assess-
ment; 22°45’N, 158°00’W) from 1988 to the present (Dore et al., 2009); (3) European Station for Time series in the Ocean (ESTOC, 29°10’N, 15°30’W) from 1994 to the present
(González-Dávila et al., 2010); and (4) Iceland Sea (IS, 68.0°N, 12.67°W) from 1985 to 2006 (Olafsson et al., 2009). Trends at the first three time-series sites are from observations
with the seasonal cycle removed. Also reported are the wintertime trends in the Iceland Sea as well as the pH difference trend for the North Pacific Ocean between transects in 1991
and 2006 (Byrne et al., 2010) and repeat sections in the western North Pacific between 1983 and 2008 (Midorikawa et al., 2010).
Notes:
* pH on the total scale.
a
Bates (2007, Table 1): Simple linear fit.
b
Bates (2007, Table 2): Seasonally detrended (including linear term for time).
c
Dore et al. (2009): Linear fit with calculated pH and pCO
2
from measured DIC and TA (full time series); corresponding Ωa from Feely et al. (2009).
d
Dore et al. (2009): Linear fit with measured pH (partial time series).
e
Santana-Casiano et al. (2007): Seasonal detrending (including linear terms for time and temperature).
f
González-Dávila et al. (2010): Seasonal detrending (including linear terms for time, temperature and mixed-layer depth).
g
Olafsson et al. (2009): Multivariable linear regression (linear terms for time and temperature) for winter data only.
h
Byrne et al. (2010): Meridional section originally occupied in 1991 and repeated in 2006.
i
Midorikawa et al. (2010): Winter and summer observations along 137°E.
j
Trends are for linear time term in seasonal detrending with harmonic periods of 12, 6 and 4 months. Harmonic analysis made after interpolating data to regular monthly grids (except for IS, which was
sampled much less frequently):
1983–2009 = September 1983 to December 2009 (BATS/Hydrostation S sampling period),
1985–2009 = February 1985 to December 2009 (IS sampling period),
1988–2009 = November 1988 to December 2009 (ALOHA/HOT sampling period), and
1995–2009 = September 1995 to December 2009 (ESTOC sampling period).
k
Ishii et al. (2011) - time-series observations in the coast of western North Pacific, with the seasonal cycle removed
l
Atmospheric pCO
2
trends computed from same harmonic analysis (12-, 6- and 4-month periods) on the GLOBALVIEW-CO
2
(2010) data product for the marine boundary layer referenced to the latitude
of the nearest atmospheric measurement station (BME = Bermuda; MLO = ALOHA; IZO = ESTOC; ICE = Iceland).
m
Winter ocean data, collected during dark period (between 19 January and 7 March), as per Olafsson et al. (2009) to reduce scatter from large interannual variations in intense short-term bloom events,
undersampled in time, fit linearly (y = at + bT + c).
300
Chapter 3 Observations: Ocean
3
cation of coastal waters (Diaz and Rosenberg, 2008; Seitzinger et al.,
2010; Kim et al., 2011), which amplifies the drawdown of CO
2
(Borges
and Gypens, 2010; Provoost et al., 2010). In addition, atmospheric
deposition of anthropogenic fixed nitrogen may now account for up
to about 3% of oceanic new production, and this nutrient source is
projected to increase (Duce et al., 2008).
Satellite observations of chlorophyll reveal that oligotrophic provinc-
es in four of the world’s major oceans expanded at average rates of
0.8 to 4.3% yr
–1
from 1998 to 2006 (Polovina et al., 2008; Irwin and
Oliver, 2009), consistent with a reduction in nutrient availability owing
to increases in stratification. Model and observational studies suggest
interannual and multi-decadal fluctuations in nutrients are coupled
with variability of mode water and the NAO in the Atlantic Ocean
(Cianca et al., 2007; Pérez et al., 2010), climate modes of variability
in the Pacific Ocean (Wong et al., 2007; Di Lorenzo et al., 2009), and
variability of subtropical gyre circulation in the Indian Ocean (Álvarez
et al., 2011). However, there are no published studies quantifying long-
term trends in ocean nutrient concentrations.
3.8.5 Conclusions
Based on high agreement between independent estimates using differ-
ent methods and data sets (e.g., oceanic carbon, oxygen, and transient
60°W 0° 60°E 120°E180°W 120°W
200 (dbar)
0°
0
100
200
300
(µmol kg
-1
)
15°N
30°N
15°S
30°S
0°
15°N
30°N
15°S
30°S
200 (dbar)
0°
15°N
30°N
15°S
30°S
200-700 (dbar)
60°W
60°W 0° 60°E 120°E180°W 120°W60°W
-16
-8
0
8
16
(µmol kg
-1
)
(a)
(b)
(c)
Figure 3.20 | Dissolved oxygen (DO) distributions (in µmol kg
–1
) between 40°S and 40°N for: (a) the climatological mean (Boyer et al., 2006) at 200 dbar, as well as changes
between 1960 and 1974 and 1990 and 2008 of (b) dissolved oxygen (∆DO) at 200 dbar and (c) ∆DO vertically averaged over 200 to 700 dbar. In (b) and (c) increases are
red and decreases blue, and areas with differences below the 95% confidence interval are shaded by black horizontal lines. (After Stramma et al., 2010.)
tracer data), it isvery likelythat the global ocean inventory of anthropo-
genic carbon (C
ant
) increased from 1994 to 2010. The oceanic C
ant
inven-
tory in 2010 is estimated to be 155 PgC with an uncertainty of ±20%.
The annual global oceanic uptake rates calculated from independent
data sets (from oceanic C
ant
inventory changes, from atmospheric O
2
/
N
2
measurements or from pCO
2
data) and for different time periods
agree with each other within their uncertainties and very likely are in
the range of 1.0 to 3.2 PgC yr
–1
.
(Section 3.8.1, Figures 3.16 and 3.17)
Oceanic uptake of anthropogenic CO
2
results in gradual acidification
of the ocean. The pH of surface seawater has decreased by 0.1 since
the beginning of the industrial era, corresponding to a 26% increase
in hydrogen ion concentration. The observed pH trends range between
–0.0014 and –0.0024 yr
–1
in surface waters. In the ocean interior, nat-
ural physical and biological processes, as well as uptake of anthro-
pogenic CO
2
, can cause changes in pH over decadal and longer time
scales (Section 3.8.2, Table 3.2, Box 3.2, Figures 3.18 and 3.19, FAQ
3.3).
High agreement among analyses provides medium confidence that
oxygen concentrations have decreased in the open ocean thermocline
in many ocean regions since the 1960s. The general decline is consist-
ent with the expectation that warming-induced stratification leads to a
decrease in the supply of oxygen to the thermocline from near surface
301
3
Observations: Ocean Chapter 3
waters, that warmer waters can hold less oxygen, and that changes in
wind-driven circulation affect oxygen concentrations. It is likely that
the tropical oxygen minimum zones have expanded in recent decades
(Section 3.8.3, Figure 3.20).
3.9 Synthesis
Substantial progress has been made since AR4 in documenting and
understanding change in the ocean. The major findings of this chapter
are largely consistent with those of AR4, but in many cases statements
can now be made with greater confidence because more data are
available, biases in historical data have been identified and reduced,
and new analytical approaches have been applied.
Changes have been observed in a number of ocean properties of
relevance to climate. It is virtually certain that the upper ocean (0 to
700 m) has warmed from 1971 to 2010 (Section 3.2.2, Figures 3.1 and
3.2). Warming between 700 and 2000 m likely contributed about 30%
of the total increase in global ocean heat content between 1957 and
2009 (Section 3.2.4, Figure 3.2). Global mean sea level has risen by
0.19 [0.17 to 0.21] m over the period 1901–2010. It is very likely that
the mean rate was 1.7 [1.5 to 1.9] mm yr
–1
between 1901 and 2010
and increased to 3.2 [2.8 to 3.6] mm yr
–1
between 1993 and 2010
(Section 3.7, Figure 3.13). The rise in mean sea level can explain most
of the observed increase in extreme sea levels (Figure 3.15). Regional
trends in sea surface salinity have very likely enhanced the mean
geographical contrasts in sea surface salinity since the 1950s: saline
surface waters in evaporation-dominated regions have become more
saline, while fresh surface waters in rainfall-dominated regions have
become fresher. It is very likely that trends in salinity have also occurred
in the ocean interior. These salinity changes provide indirect evidence
that the pattern of evaporation minus precipitation over the oceans
has been enhanced since the 1950s (Section 3.4, Figures 3.4 and 3.5].
Observed changes in water mass properties likely reflect the combined
effect of long-term trends in surface forcing (e.g., warming and
changes in evaporation minus precipitation) and variability associated
with climate modes (Section 3.5, Figure 3.9). It is virtually certain that
the ocean is storing anthropogenic CO
2
and very likely that the ocean
inventory of anthropogenic CO
2
increased from 1994 to 2010 (Section
3.8, Figures 3.16 and 3.17). The uptake of anthropogenic CO
2
has very
likely caused acidification of the ocean (Section 3.8.2, Box 3.2).
For some ocean properties, the short and incomplete observational
record is not sufficient to detect trends. For example, there is no obser-
vational evidence for or against a change in the strength of the AMOC
(Section 3.6, Figure 3.11). However, recent observations have strength-
ened evidence for variability in major ocean circulation systems and
water mass properties on time scales from years to decades. Much of
the variability observed in ocean currents and in water masses can be
linked to changes in surface forcing, including wind changes associat-
ed with the major modes of climate variability such as the NAO, SAM,
ENSO, PDO and the AMO (Section 3.6, Box 2.5).
The consistency between the patterns of change in a number of
independent ocean parameters enhances confidence in the assessment
that the physical and biogeochemical state of the oceans has changed.
This consistency is illustrated here with two simple figures (Figures 3.21
and 3.22). Four global measures of ocean change have increased since
the 1950s: the inventory of anthropogenic CO
2
, global mean sea level,
upper ocean heat content, and the salinity contrast between regions of
high and low sea surface salinity (Figure 3.21). High agreement among
multiple lines of evidence based on independent data and different
methods provides high confidence in the observed increase in these
global metrics of ocean change.
The distributions of trends in subsurface water properties, summarized
in a schematic zonally averaged view in Figure 3.22, are consistent
both with each other and with well-understood dynamics of ocean
circulation and water mass formation. The largest changes in
temperature, salinity, anthropogenic CO
2
, and other properties are
observed along known ventilation pathways (indicated by arrows in
Figure 3.22), where surface waters are transferred to the ocean interior,
or in regions where changes in ocean circulation (e.g., contraction or
expansion of gyres, or a southward shift of the Antarctic Circumpolar
1950 1960 1970 1980 1990 2000 2010
0
50
100
150
(mm)
Global mean sea level
1950 1960 1970 1980 1990 2000 2010
0
50
100
150
200
250
(ZJ)
Upper ocean heat content
1950 1960 1970 1980 1990 2000 2010
-0.09
-0.06
-0.03
0
0.03
0.06
0.09
(PSS-78)
High salinity minus low salinity
Year
1950 1960 1970 1980 1990 2000 2010
0
25
50
75
100
125
(PgC)
Carbon
(a)
(b)
(c)
(d)
Figure 3.21 | Time series of changes in large-scale ocean climate properties. From
top to bottom: global ocean inventory of anthropogenic carbon dioxide, updated from
Khatiwala et al. (2009); global mean sea level (GMSL), from Church and White (2011);
global upper ocean heat content anomaly, updated from Domingues et al. (2008); the
difference between salinity averaged over regions where the sea surface salinity is
greater than the global mean sea surface salinity (“High Salinity”) and salinity averaged
over regions values below the global mean (“Low Salinity”), from Boyer et al. (2009).
302
Chapter 3 Observations: Ocean
3
Current) result in large anomalies. Zonally averaged warming trends
are widespread throughout the upper 2000 m, with largest warming
near the sea surface. Water masses formed in the precipitation-
dominated mid to high latitudes have freshened, while water masses
formed in the evaporation-dominated subtropics have become saltier.
Anthropogenic CO
2
has accumulated in surface waters and been
transferred into the interior, primarily by water masses formed in the
North Atlantic and Southern Oceans.
In summary, changes have been observed in ocean properties of rele-
vance to climate during the past 40 years, including temperature, salin-
ity, sea level, carbon, pH, and oxygen. The observed patterns of change
are consistent with changes in the surface ocean (warming, changes
in salinity and an increase in C
ant
) in response to climate change and
variability and with known physical and biogeochemical processes in
the ocean, providing high confidence in this assessment. Chapter 10
discusses the extent to which these observed changes can be attribut-
ed to human or natural forcing.
Improvements in the quality and quantity of ocean observations has
allowed for a more definitive assessment of ocean change than was
possible in AR4. However, substantial uncertainties remain. In many
cases, the observational record is still too short or incomplete to detect
trends in the presence of energetic variability on time scales of years
to decades. Recent improvements in the ocean observing system, most
notably the Argo profiling float array, mean that temperature and
salinity are now being sampled routinely in most of the ocean above
2000 m depth for the first time. However, sparse sampling of the deep
ocean and of many biogeochemical variables continues to limit the
ability to detect and understand changes in the global ocean.
Depth (m)
Temperature change (°C per decade)
Figure 3.22 | Summary of observed changes in zonal averages of global ocean properties. Temperature trends (degrees Celsius per decade) are indicated in colour (red = warm-
ing, blue = cooling); salinity trends are indicated by contour lines (dashed = fresher; solid = saltier) for the upper 2000 m of the water column (50-year trends from data set of
Durack and Wijffels (2010); trends significant at >90% confidence are shown). Arrows indicate primary ventilation pathways. Changes in other physical and chemical properties are
summarised to the right of the figure, for each depth range (broken axes symbols delimit changes in vertical scale). Increases are shown in red, followed by a plus sign; decreases
are shown in blue, followed by a minus sign; the number of + and – signs indicates the level of confidence associated with the observation of change (+++, high confidence; ++,
medium confidence; +, low confidence). T = temperature, S = salinity, Strat = stratification, C
ant
= anthropogenic carbon, CO
3
2–
= carbonate ion, NA = North Atlantic, SO = Southern
Ocean, AABW = Antarctic Bottom Water. S > S
–
refers to the salinity averaged over regions where the sea surface salinity is greater than the global mean sea surface salinity; S <
S
–
refers to the average over regions with values below the global mean.
303
3
Observations: Ocean Chapter 3
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Observations: Ocean Chapter 3
Appendix 3.A:
Availability of Observations for Assessment of
Change in the Oceans
Sampling of the ocean has been highly heterogeneous since 1950.
The coverage in space, time, depth and number of ocean variables
has evolved over time, reflecting changes in technology and the con-
tribution of major oceanographic research programs. Changes in the
distribution and quality of ocean measurements over time complicate
efforts to detect and interpret change in the ocean. This Appendix pro-
vides some illustrative examples of the evolution of the ocean observ-
ing system on which the assessment of ocean change in this chapter
is based. A more comprehensive discussion of ocean sampling is pro-
vided in the literature cited in this chapter. Sampling of sea surface
temperature is discussed in Chapter 2.
3.A.1 Subsurface Ocean Temperature and Heat Content
Temperature is the best-sampled oceanographic variable, but even
for temperature sampling is far from ideal or complete. Early oceano-
graphic expeditions included the Challenger voyage around the world
in the 1870s, the Meteor survey of the Atlantic in the 1920s, and the
Discovery investigations of the Southern Ocean starting in the 1920s.
More frequent basin-scale sampling commenced in the late 1950s with
the International Geophysical Year. The number of profiles available for
assessment of changes in temperature and ocean heat content in the
upper 700 m generally increases with time since the 1950s (Figure
3.A.1). Near-global coverage of the upper half of the ocean was not
achieved until the widespread deployment of Argo profiling floats in
the 2000s (Figure 3.A.2).
1950s 1960s 1970s
60°S
30°S
0°
30°N
60°N
1980s
60°E 160°W 20°W
2000s
60°E 160°W 20°W
60°S
30°S
0°
30°N
60°N
5
1
5
2
5
1990s
No. Temperature Profiles 0−700 m ( 1° x 1° )
60°E 160°W 20°W
Figure 3.A.1 | Number of temperature profiles extending to 700 m depth in each 1° × 1° square, by decade, between 65°N and 65°S.
Early measurements of temperature were made using reversing ther-
mometers and Nansen bottles that were lowered from ships on station
(not moving). Starting in the 1960s conductivity-temperature-depth
(CTD) instruments with Niskin bottles gradually gained dominance for
high-quality data and deep data collected on station during oceano-
graphic cruises. From at least 1950 through about 1970, most subsur-
face measurements of ocean temperature were made with mechanical
bathythermographs, an advance because these instruments could be
deployed from a moving ship, albeit a slowly moving one, but these
casts were generally limited to depths shallower than 250 m. Expend-
able bathythermographs (XBTs) that could be deployed from a rapidly
moving ship and sampled to 400 m came into widespread use in the
late 1960s, and those that sampled to 700 m became predominant
in the 1990s, greatly expanding oceanographic sampling. Starting in
2000, Argo floats began sampling the ocean to a target depth of 2000
m, building to near-global coverage by 2005. Prior to the Argo era,
sampling of the ocean below 700 m was almost solely achieved from
ships on station deploying Nansen bottles with reversing thermom-
eters or later using CTDs with Niskin bottles. Today ship-based station
data still dominates sampling for waters deeper than 2000 m depth
(the maximum depth currently sampled by Argo floats). An illustration
of the limited data available for assessment of change in the deep
ocean is provided in Figure 3.3, which shows locations of full-depth
oceanographic CTD sections that have been occupied more than once
since about 1980. The depth coverage of the ocean observing system
has changed over time (Figure 3A.2, top panel) with a hemispheric bias
(Figure 3.A.2, middle and lower panels). The Northern Hemisphere (NH)
has been consistently better sampled than the Southern Hemisphere
(SH) prior to the Argo era.
312
Chapter 3 Observations: Ocean
3
Figure 3.A.2 | (Top) Percentage of global coverage of ocean temperature profiles as a
function of depth in 1° latitude by 1° longitude by 1-year bins (top panel) shown versus
time. Different colours indicate profiles to different depths (middle panel). Percentage of
global coverage as a function of depth and time, for the Northern Hemisphere. (Bottom
panel) As above, but for the Southern Hemisphere.
1950 1960 1970 1980 1990 2000 2010
0
20
40
60
80
Ocean temperature profiles − Yearly coverage
0−100 m
0−200 m
0−300 m
0−400 m
0−700 m
0−900 m
0−1500 m
0−1800 m
GLOBAL
Global coverage (%)
NORTHERN HEMISPHERE
Depth (m)
1950 1960 1970 1980 1990 2000 2010
0
500
1000
1500
2000
SOUTHERN HEMISPHERE
Depth (m)
1950 1960 1970 1980 1990 2000 2010
0
500
1000
1500
2000
Global coverage (%)
0 10 20 30 40 50 60 70
3.A.2 Salinity
Measurements of subsurface salinity have relied almost solely on data
collected from bottle and CTD casts from ships on station (and, more
recently, using profiling floats that sample both temperature and salin-
ity with CTDs). Hence fewer measurements of salinity are available
than of temperature (by roughly a factor of 3). However, the evolution
with time of subsurface salinity sampling shows a progression simi-
lar to that of temperature (Figure 3.A.3). Coverage generally improves
with time, but there is a strong NH bias, particularly in the North
Atlantic. A shift in focus from basin-to-basin as major field programs
were carried out is evident. Near-global coverage of ocean salinity
above 2000 m was not achieved until after 2005, when the Argo array
approached full deployment. For depths greater than 2000 m, outside
the relatively well sampled North Atlantic, information on changes in
ocean salinity is largely restricted to the repeat hydrographic transects
(see Figure 3.3).
3.A.3 Sea Level
Direct observations of sea level are made using tide gauges since the
1700s and high-precision satellite altimeters since 1992. Tide gauge
measurements are limited to coastlines and islands. There are intermit-
tent records of sea level at Amsterdam from 1700 and at three more
sites in Northern Europe starting after 1770. By the late 1800s, more
tide gauges were being operated in Northern Europe and in North
America, as well as in Australia and New Zealand (Figure 3.A.4). It
was not until the late 1970s to early 1980s that a majority of deep-
ocean islands had operating tide gauges suitable for climate studies.
Although tide gauges have continued to be deployed since 1990, they
have been complemented by continuous, near-global measurements
of sea level from space since 1992. Measurements are made along the
satellite’s ground track on the Earth surface, typically averaged over
approximately 7 km to reduce noise and improve precision. The maxi-
mum latitude extent of the measurement is limited by the inclination
of the orbital plane, which has been between ±66° for the TOPEX/
Poseidon and Jason series of altimeters. The spacing between ground
tracks is much greater than the spacing along the ground track. As
an example, the groundtrack separation of the TOPEX/Poseidon-type
of orbits is about 300 km at the equator, but is less than 100 km at
latitudes poleward of 50° latitude. On average, the spacing is between
100 and 200 km. Satellites are limited in the temporal sampling as
well due to the orbit configuration. For a specific location along a
groundtrack, the return time for a TOPEX/Poseidon-type of orbit is 9.9
days. If one relaxes the requirement to a measurement within a 300 km
radius, the return time can be as short as a few hours at high latitudes
to about 3 days at the equator. As noted in Section 3.6, satellite altim-
eter observations of sea level are also an important tool for observing
large-scale ocean circulation.
3.A.4 Biogeochemistry
The data available for assessing changes in biogeochemical param-
eters is less complete than for temperature and salinity. The global
data base on which the Global Ocean Data Analysis Project (GLODAP,
Key et al., 2004) ocean carbon inventory is based is illustrated in Figure
3.A.5. Changes in the ocean inventory of anthropogenic CO
2
have been
estimated using measurements of carbon parameters and other tracers
collected at these roughly 12,000 stations, mostly occupied since 1990.
The majority of these stations extend through the full water depth. A
subset of these stations have been repeated one or more times. The
distribution of oxygen measurements at 300 m depth in 10-year peri-
ods since 1960 is shown in Figure 3.A.6, as used in the global study of
Stramma et al. (2012). As for temperature and salinity, the sampling
is heterogeneous, coverage generally improves with time but shifts
between basins as major field programs come and go, and tends to be
concentrated in the NH.
313
3
Observations: Ocean Chapter 3
Latitude
1980 to 1985
Total: 152305
Atlantic: 84119
Pacific: 28252
Indian: 8359
75S
55S
35S
15S
15N
35N
55N
75N
1985 to 1990
Total: 164376
Atlantic: 94367
Pacific: 28323
Indian: 8625
1990 to 1995
Total: 126420
Atlantic: 61562
Pacific: 31443
Indian: 4195
Latitude
Longitude
1995 to 2000
Total: 63356
Atlantic: 36210
Pacific: 9123
Indian: 5614
0 60E 120E 180 120W 60W 0
75S
55S
35S
15S
15N
35N
55N
75N
Longitude
2000 to 2005
Total: 98169
Atlantic: 30290
Pacific: 46613
Indian: 18133
0 60E 120E 180 120W 60W 0
Longitude
2005 to 2010
Total: 369960
Atlantic: 74633
Pacific: 207567
Indian: 75888
0 60E 120E 180 120W 60W 0
Latitude
1950 to 1955
Total: 45556
Atlantic: 25182
Pacific: 10605
Indian: 226
75S
55S
35S
15S
15N
35N
55N
75N
1955 to 1960
Total: 58519
Atlantic: 29123
Pacific: 14385
Indian: 1184
1960 to 1965
Total: 82505
Atlantic: 38697
Pacific: 17573
Indian: 5584
Latitude
1965 to 1970
Total: 117464
Atlantic: 50564
Pacific: 29262
Indian: 6845
75S
55S
35S
15S
15N
35N
55N
75N
1970 to 1975
Total: 137089
Atlantic: 68844
Pacific: 24345
Indian: 8614
1975 to 1980
Total: 140434
Atlantic: 67843
Pacific: 27029
Indian: 10213
Figure 3.A.3 | Hydrographic profile data used in the Durack and Wijffels (2010) study. Station locations for 5-year temporal bins from 1950–1955 (top left) to
2005–2010 (bottom right).
314
Chapter 3 Observations: Ocean
3
1880-1890
1900-1910
1920-1930
1940-1950
1960-1970
1980-1990
Figure 3.A.4 | Location of tide gauges (red dots) that had at least 1 year of observations within the decade indicated.
Figure 3.A.5 | Location of profiles used to construct the Global Ocean Data Analysis Project (GLODAP) ocean carbon climatology.
60
o
S
30
o
S
0
o
30
o
N
60
o
N
GLODAP Stations (12011)
315
3
Observations: Ocean Chapter 3
180
o
W
120
o
W
60
o
W
0
o
60
o
E 120
o
E 180
o
W
60
o
S
30
o
S
0
o
30
o
N
60
o
N
1950 − 1960 1960 − 1970
1970 − 1980 1980 − 1990
1990 − 2000 2000 − 2010
180
o
W
120
o
W
60
o
W
0
o
60
o
E 120
o
E 180
o
W
180
o
W
120
o
W
60
o
W
0
o
60
o
E 120
o
E 180
o
W 180
o
W
120
o
W
60
o
W
0
o
60
o
E 120
o
E 180
o
W
180
o
W
120
o
W
60
o
W
0
o
60
o
E 120
o
E 180
o
W 180
o
W
120
o
W
60
o
W
0
o
60
o
E 120
o
E 180
o
W
60
o
S
30
o
S
0
o
30
o
N
60
o
N
60
o
S
30
o
S
0
o
30
o
N
60
o
N
60
o
S
30
o
S
0
o
30
o
N
60
o
N
60
o
S
30
o
S
0
o
30
o
N
60
o
N
60
o
S
30
o
S
0
o
30
o
N
60
o
N
Figure 3.A.6 | Distribution of oxygen measurements at 300 dbar for the decades 1950 to 1960 (upper left) to 2000 to 2010 (lower right frame). (From Stramma et al., 2012.)
[Note that additional oxygen data have become available for the 2000–2010 period since that study was completed.]
