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Observations: Oceanic Climate Change and Sea Level

Oceanic Climate Change and Sea Level
Coordinating Lead Authors:
Nathaniel L. Bindoff (Australia), Jürgen Willebrand (Germany)
Lead Authors:
Vincenzo Artale (Italy), Anny Cazenave (France), Jonathan M. Gregory (UK), Sergey Gulev (Russian Federation), Kimio Hanawa (Japan),
Corrine Le Quéré (UK, France, Canada), Sydney Levitus (USA), Yukihiro Nojiri (Japan), C.K. Shum (USA), Lynne D. Talley (USA),
Alakkat S. Unnikrishnan (India)
Contributing Authors:
J. Antonov (USA, Russian Federation), N.R. Bates (Bermuda), T. Boyer (USA), D. Chambers (USA), B. Chao (USA), J. Church (Australia),
R. Curry (USA), S. Emerson (USA), R. Feely (USA), H. Garcia (USA), M. González-Davíla (Spain), N. Gruber (USA, Switzerland),
S. Josey (UK), T. Joyce (USA), K. Kim (Republic of Korea), B. King (UK), A. Koertzinger (Germany), K. Lambeck (Australia),
K. Laval (France), N. Lefevre (France), E. Leuliette (USA), R. Marsh (UK), C. Mauritzen (Norway), M. McPhaden (USA), C. Millot (France),
C. Milly (USA), R. Molinari (USA), R.S. Nerem (USA), T. Ono (Japan), M. Pahlow (Canada), T.-H. Peng (USA), A. Proshutinsky (USA),
B. Qiu (USA), D. Quadfasel (Germany), S. Rahmstorf (Germany), S. Rintoul (Australia), M. Rixen (NATO, Belgium), P. Rizzoli (USA, Italy),
C. Sabine (USA), D. Sahagian (USA), F. Schott (Germany), Y. Song (USA), D. Stammer (Germany), T. Suga (Japan), C. Sweeney (USA),
M. Tamisiea (USA), M. Tsimplis (UK, Greece), R. Wanninkhof (USA), J. Willis (USA), A.P.S. Wong (USA, Australia), P. Woodworth (UK),
I. Yashayaev (Canada), I. Yasuda (Japan)
Review Editors:
Laurent Labeyrie (France), David Wratt (New Zealand)
This chapter should be cited as:
Bindoff, N.L., J. Willebrand, V. Artale, A, Cazenave, J. Gregory, S. Gulev, K. Hanawa, C. Le Quéré, S. Levitus, Y. Nojiri, C.K. Shum, L.D.
Talley and A. Unnikrishnan, 2007: Observations: Oceanic Climate Change and Sea Level. In: Climate Change 2007: The Physical Science
Basis. Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change [Solomon, S.,
D. Qin, M. Manning, Z. Chen, M. Marquis, K.B. Averyt, M. Tignor and H.L. Miller (eds.)]. Cambridge University Press, Cambridge, United
Kingdom and New York, NY, USA.
Observations: Oceanic Climate Change and Sea Level Chapter 5
Table of Contents
Executive Summary .................................................... 387
5.1 Introduction ......................................................... 389
5.2 Changes in Global-Scale Temperature
and Salinity .......................................................... 389
5.2.1 Background ......................................................... 389
5.2.2 Ocean Heat Content ............................................ 390
5.2.3 Ocean Salinity ......................................................393
5.2.4 Air-Sea Fluxes and Meridional Transports ........... 393
5.3 Regional Changes in Ocean Circulation
and Water Masses .............................................. 394
5.3.1 Introduction .........................................................394
5.3.2 Atlantic and Arctic Oceans .................................. 395
Box 5.1: Has the Meridional Overturning Circulation
in the Atlantic Changed? ..................................... 397
5.3.3 Pacific Ocean ...................................................... 399
5.3.4 Indian Ocean ....................................................... 400
5.3.5 Southern Ocean .................................................. 401
5.3.6 Relation of Regional to Global Changes ............. 402
5.4 Ocean Biogeochemical Changes ................... 403
5.4.1 Introduction .........................................................403
5.4.2 Carbon ................................................................. 403
5.4.3 Oxygen ................................................................406
5.4.4 Nutrients .............................................................. 406
5.4.5 Biological Changes Relevant to Ocean
Biogeochemistry .................................................. 408
5.4.6 Consistency with Physical Changes ................... 408
5.5 Changes in Sea Level ......................................... 408
5.5.1 Introductory Remarks .......................................... 408
5.5.2 Observations of Sea Level Changes ...................410
5.5.3 Ocean Density Changes ...................................... 414
5.5.4 Interpretation of Regional Variations in the
Rate of Sea Level Change ................................... 416
5.5.5 Ocean Mass Change ........................................... 417
5.5.6 Total Budget of the Global Mean Sea
Level Change ....................................................... 419
5.6 Synthesis ............................................................... 420
Frequently Asked Question
FAQ 5.1: Is Sea Level Rising? ............................................409
References ........................................................................ 422
Appendix 5.A: Techniques, Error
Estimation and Measurement Systems ................. 429
Chapter 5 Observations: Oceanic Climate Change and Sea Level
Executive Summary
The oceans are warming. Over the period 1961 to 2003,
global ocean temperature has risen by 0.10°C from the
surface to a depth of 700 m. Consistent with the Third
Assessment Report (TAR), global ocean heat content (0–
3,000 m) has increased during the same period, equivalent
to absorbing energy at a rate of 0.21 ± 0.04 W m–2 globally
averaged over the Earth’s surface. Two-thirds of this energy
is absorbed between the surface and a depth of 700 m.
Global ocean heat content observations show considerable
interannual and inter-decadal variability superimposed on
the longer-term trend. Relative to 1961 to 2003, the period
1993 to 2003 has high rates of warming but since 2003
there has been some cooling.
Large-scale, coherent trends of salinity are observed for
1955 to 1998, and are characterised by a global freshening
in subpolar latitudes and a salinication of shallower
parts of the tropical and subtropical oceans. Freshening is
pronounced in the Pacic while increasing salinities prevail
over most of Atlantic and Indian Oceans. These trends are
consistent with changes in precipitation and inferred larger
water transport in the atmosphere from low latitudes to high
latitudes and from the Atlantic to the Pacic. Observations
do not allow for a reliable estimate of the global average
change in salinity in the oceans.
Key oceanic water masses are changing; however, there is
no clear evidence for ocean circulation changes. Southern
Ocean mode waters and Upper Circumpolar Deep Waters
have warmed from the 1960s to about 2000. A similar
but weaker pattern of warming in the Gulf Stream and
Kuroshio mode waters in the North Atlantic and North
Pacic has been observed. Long-term cooling is observed
in the North Atlantic subpolar gyre and in the central North
Pacic. Since 1995, the upper North Atlantic subpolar gyre
has been warming and becoming more saline. It is very
likely that up to the end of the 20th century, the Atlantic
meridional overturning circulation has been changing
signicantly at interannual to decadal time scales. Over the
last 50 years, no coherent evidence for a trend in the strength
of the meridional overturning circulation has been found.
Ocean biogeochemistry is changing. The total inorganic
carbon content of the oceans has increased by 118 ±
19 GtC between the end of the pre-industrial period (about
1750) and 1994 and continues to increase. It is more likely
than not that the fraction of emitted carbon dioxide that was
taken up by the oceans has decreased, from 42 ± 7% during
1750 to 1994 to 37 ± 7% during 1980 to 2005. This would
be consistent with the expected rate at which the oceans
can absorb carbon, but the uncertainty in this estimate does
not allow rm conclusions. The increase in total inorganic
carbon caused a decrease in the depth at which calcium
carbonate dissolves, and also caused a decrease in surface
ocean pH by an average of 0.1 units since 1750. Direct
observations of pH at available time series stations for the
last 20 years also show trends of decreasing pH at a rate of
0.02 pH units per decade. There is evidence for decreased
oxygen concentrations, likely driven by reduced rates of
water renewal, in the thermocline (~100–1,000 m) in most
ocean basins from the early 1970s to the late 1990s.
Global mean sea level has been rising. From 1961 to
2003, the average rate of sea level rise was 1.8 ± 0.5
mm yr–1. For the 20th century, the average rate was 1.7 ± 0.5
mm yr–1, consistent with the TAR estimate of 1 to 2 mm yr–1.
There is high condence that the rate of sea level rise
has increased between the mid-19th and the mid-20th
centuries. Sea level change is highly non-uniform spatially,
and in some regions, rates are up to several times the
global mean rise, while in other regions sea level is falling.
There is evidence for an increase in the occurrence of
extreme high water worldwide related to storm surges, and
variations in extremes during this period are related to the
rise in mean sea level and variations in regional climate.
The rise in global mean sea level is accompanied by
considerable decadal variability. For the period 1993 to
2003, the rate of sea level rise is estimated from observations
with satellite altimetry as 3.1 ± 0.7 mm yr–1, signicantly
higher than the average rate. The tide gauge record indicates
that similar large rates have occurred in previous 10-year
periods since 1950. It is unknown whether the higher rate
in 1993 to 2003 is due to decadal variability or an increase
in the longer-term trend.
There are uncertainties in the estimates of the contributions
to sea level change but understanding has signicantly
improved for recent periods. For the period 1961 to 2003,
the average contribution of thermal expansion to sea level
rise was 0.4 ± 0.1 mm yr–1. As reported in the TAR, it is
likely that the sum of all known contributions for this period
is smaller than the observed sea level rise, and therefore it
is not possible to satisfactorily account for the processes
causing sea level rise. However, for the period 1993 to
2003, for which the observing system is much better, the
contributions from thermal expansion (1.6 ± 0.5 mm yr–1)
and loss of mass from glaciers, ice caps and the Greenland
and Antarctic Ice Sheets together give 2.8 ± 0.7 mm yr–1.
For the latter period, the climate contributions constitute
the main factors in the sea level budget, which is closed to
within known errors.
Observations: Oceanic Climate Change and Sea Level Chapter 5
The patterns of observed changes in global ocean heat
content and salinity, sea level, thermal expansion, water
mass evolution and biogeochemical parameters described
in this chapter are broadly consistent with the observed
ocean surface changes and the known characteristics of the
large-scale ocean circulation.
Chapter 5 Observations: Oceanic Climate Change and Sea Level
5.1 Introduction
The ocean has an important role in climate variability
and change. The ocean’s heat capacity is about 1,000 times
larger than that of the atmosphere, and the oceans net heat
uptake since 1960 is around 20 times greater than that of the
atmosphere (Levitus et al., 2005a). This large amount of heat,
which has been mainly stored in the upper layers of the ocean,
plays a crucial role in climate change, in particular variations
on seasonal to decadal time scales. The transport of heat and
freshwater by ocean currents can have an important effect on
regional climates, and the large-scale Meridional Overturning
Circulation (MOC; also referred to as thermohaline circulation)
inuences the climate on a global scale (e.g., Vellinga and
Wood, 2002). Life in the sea is dependent on the biogeochemical
status of the ocean and is inuenced by changes in the physical
state and circulation. Changes in ocean biogeochemistry can
directly feed back to the climate system, for example, through
changes in the uptake or release of radiatively active gases such
as carbon dioxide. Changes in sea level are also important for
human society, and are linked to changes in ocean circulation.
Finally, oceanic parameters can be useful for detecting climate
change, in particular temperature and salinity changes in the
deeper layers and in different regions where the short-term
variability is smaller and the signal-to-noise ratio is higher.
The large-scale, three-dimensional ocean circulation and the
formation of water masses that ventilate the main thermocline
together create pathways for the transport of heat, freshwater
and dissolved gases such as carbon dioxide from the surface
ocean into the density-stratied deeper ocean, thereby isolating
them from further interaction with the atmosphere. These
pathways are also important for the transport of anomalies in
these parameters caused by changes in the surface conditions.
Furthermore, changes in the storage of heat and in the distribution
of ocean salinity cause the ocean to expand or contract and
hence change the sea level both regionally and globally.
The ocean varies over a broad range of time scales, from
seasonal (e.g., in the surface mixed layer) to decadal (e.g.,
circulation in the main subtropical gyres) to centennial and
longer (associated with the MOC). The main modes of climate
variability, which are described in Chapter 3, are the El Niño-
Southern Oscillation (ENSO), the Pacic Decadal Oscillation
(PDO), the Northern Annular Mode (NAM), which is related
to the North Atlantic Oscillation (NAO), and the Southern
Annular Mode (SAM). Forcing of the oceans is often related to
these modes, which cause changes in ocean circulation through
changed patterns of winds and changes in surface ocean
The Third Assessment Report (TAR) discussed some aspects
of the ocean’s role. Folland et al. (2001) concluded that the
global ocean has signicantly warmed since the late 1950’s.
This assessment provides updated estimates of temperature
changes for the oceans. Furthermore, it discusses new evidence
for changes in the ocean freshwater budget and the ocean
circulation. The TAR estimate of the total inorganic carbon
increase in the ocean (Prentice et al., 2001) was based entirely
on indirect evidence. This assessment provides updated indirect
estimates and reports on new and direct evidence for changes in
total carbon increase and for changes in ocean biogeochemistry
(including pH and oxygen). Church et al. (2001) determined a
range of 1 to 2 mm yr–1 for the observed global average sea level
rise in the 20th century. This assessment provides new estimates
for sea level change and the climate-related contributions to sea
level change from thermal expansion and melting of ice sheets,
glaciers and ice caps. The focus of this chapter is on observed
changes in the global ocean basins, however some regional
changes in the ocean state are also considered.
Many ocean observations are poorly sampled in space and
time, and regional distributions often are quite heterogeneous.
Furthermore, the observational records only cover a relatively
short period of time (e.g., the 1950s to the present). Many of the
observed changes have signicant decadal variability associated
with them, and in some cases decadal variability and/or poor
sampling may prevent detection of long-term trends. When
time series of oceanic parameters are considered, linear trends
are often computed in order to quantify the observed long-term
changes; however, this does not imply that the original signal
is best represented by a linear increase in time. For plotting
time series, this chapter generally uses the difference (anomaly)
from the average value for the years 1961 to 1990. Wherever
possible, error bars are provided to quantify the uncertainty of
the observations. As in other parts of this report, 90% condence
intervals are used throughout. If not otherwise stated, values
with error bars given as x ± e should hence be interpreted as a
90% chance that the true value is in the range x – e to x + e.
5.2 Changes in Global-Scale
Temperature and Salinity
5.2.1 Background
Among the major challenges in understanding the climate
system are quantifying the Earth’s heat balance and the
freshwater balance (hydrological cycle), which both have a
substantial contribution from the World Ocean. This chapter
presents observational evidence that directly or indirectly helps
to quantify changes in these balances.
The TAR included estimates of ocean heat content changes
for the upper 3,000 m of the World Ocean. Ocean heat content
change is closely proportional to the average temperature
change in a volume of seawater, and is dened here as the
deviation from a reference period. This section reports on
updates of this estimate and presents estimates for the upper
700 m based on additional modern and historical data (Willis et
al., 2004; Levitus et al., 2005b; Ishii et al., 2006). The section
also presents new estimates of the temporal variability of salinity.
The data used for temperature and heat content estimates are
based on the World Ocean Database 2001 (e.g., Boyer et al.,
Observations: Oceanic Climate Change and Sea Level Chapter 5
2002; Conkright et al., 2002), which has been updated with
more recent data. Temperature data include measurements
from reversing thermometers, expendable bathythermographs,
mechanical bathythermographs, conductivity-temperature-
depth instruments, Argo proling oats, moored buoys and
drifting buoys. The salinity data are described by Locarnini et
al. (2002) and Stephens et al. (2002).
5.2.2 Ocean Heat Content Long-Term Temperature Changes
Figure 5.1 shows two time series of ocean heat content for
the 0 to 700 m layer of the World Ocean, updated from Ishi et
al. (2006) and Levitus et al. (2005a) for 1955 to 2005, and a
time series for 0 to 750 m for 1993 to 2005 updated from Willis
et al. (2004). Approximately 7.9 million temperature proles
were used in constructing the two longer time series. The three
heat content analyses cover different periods but where they
overlap in time there is good agreement. The time series shows
an overall trend of increasing heat content in the World Ocean
with interannual and inter-decadal variations superimposed on
this trend. The root mean square difference between the three
data sets is 1.5 × 1022 J. These year-to-year differences, which
are due to differences in quality control and data used, are
small and now approaching the accuracies required to close the
Earth’s radiation budget (e.g., Carton et al., 2005). On longer
time scales, the two longest time series (using independent
criteria for selection, quality control, interpolation and analysis
of similar data sets) show good agreement about long-term
trends and also on decadal time scales.
For the period 1993 to 2003, the Levitus et al. (2005a)
analysis has a linear global ocean trend of 0.42 ± 0.18 W m–2,
Willis et al. (2004) has a trend of 0.66 ± 0.18 W m–2 and Ishii
et al. (2006) a trend of 0.33 ± 0.18 W m–2. Overall, we assess
the trend for this period as 0.5 ± 0.18 W m–2. For the 0 to 700
m layer and the period 1955 to 2003 the heat content change
is 10.9 ± 3.1 × 1022 J or 0.14 ± 0.04 W m–2 (data from Levitus
et al., 2005a). All of these estimates are per unit area of Earth
surface. Despite the fact that there are differences between these
three ocean heat content estimates due to the data used, quality
control applied, instrumental biases, temporal and spatial
averaging and analysis methods (Appendix 5.A.1), they are
consistent with each other giving a high degree of condence
for their use in climate change studies. The global increase
in ocean heat content during the period 1993 to 2003 in two
ocean models constrained by assimilating altimetric sea level
and other observations (Carton et al., 2005; Köhl et al., 2006)
is considerably larger than these observational estimates. We
assess the heat content change from both of the long time series
(0 to 700 m layer and the 1961 to 2003 period) to be 8.11 ±
0.74 × 1022 J, corresponding to an average warming of 0.1°C or
0.14 ± 0.04 W m–2, and conclude that the available heat content
estimates from 1961 to 2003 show a signicant increasing trend
in ocean heat content.
The data used in estimating the Levitus et al. (2005a)
ocean temperature elds (for the above heat content estimates)
do not include sea surface temperature (SST) observations,
Figure 5.1. Time series of global annual ocean heat content (1022 J) for the 0 to 700 m layer. The black curve is updated from Levitus et al. (2005a), with the shading repre-
senting the 90% confidence interval. The red and green curves are updates of the analyses by Ishii et al. (2006) and Willis et al. (2004, over 0 to 750 m) respectively, with the er-
ror bars denoting the 90% confidence interval. The black and red curves denote the deviation from the 1961 to 1990 average and the shorter green curve denotes the deviation
from the average of the black curve for the period 1993 to 2003.
Chapter 5 Observations: Oceanic Climate Change and Sea Level
which are discussed in Chapter 3. However, comparison of the
global, annual mean time series of near-surface temperature
(approximately 0 to 5 m depth) from this analysis and the
corresponding SST series based on a subset of the International
Comprehensive Ocean-Atmosphere Data Set (ICOADS)
database (approximately 134 million SST observations; Smith
and Reynolds, 2003 and additional data) shows a high correlation
(r = 0.96) for the period 1955 to 2005. The consistency between
these two data sets gives condence in the ocean temperature
data set used for estimating depth-integrated heat content, and
supports the trends in SST reported in Chapter 3.
There is a contribution to the global heat content integral
from depths greater than 700 m as documented by Levitus et al.
(2000; 2005a). However, due to the lack of data with increasing
depth the data must be composited using ve-year running
pentads in order to have enough data for a meaningful analysis
in the deep ocean. Even then, there are not enough deep ocean
data to extend the time series for the upper 3,000 m past the
1994–1998 pentad. There is a close correlation between the 0
to 700 and 0 to 3,000 m time series of Levitus et al. (2005a). A
comparison of the linear trends from these two series indicates
that about 69% of the increase in ocean heat content during
1955 to 1998 (the period when estimates from both time series
are available) occurred in the upper 700 m of the World Ocean.
Based on the linear trend, for the 0 to 3,000 m layer for the period
1961 to 2003 there has been an increase of ocean heat content
of approximately 14.2 ± 2.4 × 1022 J, corresponding to a global
ocean volume mean temperature increase of 0.037°C during
this period. This increase in ocean heat content corresponds to
an average heating rate of 0.21 ± 0.04 W m–2 for the Earth’s
The geographical distribution of the linear trend of 0 to
700 m heat content for 1955 to 2003 for the World Ocean is
shown in Figure 5.2. These trends are non-uniform in space,
with some regions showing cooling and others warming. Most
of the Atlantic Ocean exhibits warming with a major exception
being the subarctic gyre. The Atlantic Ocean accounts for
approximately half of the global linear trend of ocean heat
content (Levitus et al., 2005a). Much of the Indian Ocean has
warmed since 1955 with a major exception being the 5°S to
20°S latitude belt. The Southern Ocean (south of 35°S) in the
Atlantic, Indian and Pacic sectors has generally warmed.
The Pacic Ocean is characterised by warming with major
exceptions along 40°N and the western tropical Pacic.
Figure 5.3 shows the linear trends (1955 to 2003) of zonally
averaged temperature anomalies (0 to 1,500 m) for the World
Ocean and individual basins based on yearly anomaly elds
(Levitus et al., 2005a). The strongest trends in these anomalies
are concentrated in the upper ocean. Warming occurs at most
latitudes in all three of the ocean basins. The regions that
exhibit cooling are mainly in the shallow equatorial areas and in
some high-latitude regions. In the Indian Ocean, cooling occurs
at subsurface depths centred on 12°S at 150 m depth and in
the Pacic centred on the equator and 150 m depth. Cooling
also occured in the 32°N to 48°N region of the Pacic Ocean
and the 49°N to 60°N region of the Atlantic Ocean. Regional
temperature changes are discussed further in Section 5.3. Variability of Heat Content
A major feature of Figure 5.1 is the relatively large increase
in global ocean heat content during 1969 to 1980 and a sharp
Figure 5.2. Linear trends (1955–2003) of change in ocean heat content per unit surface area (W m–2) for the 0 to 700 m layer, based on the work of Levitus et al. (2005a). The
linear trend is computed at each grid point using a least squares fit to the time series at each grid point. The contour interval is 0.25 W m–2. Red shading indicates values equal
to or greater than 0.25 W m–2 and blue shading indicates values equal to or less than –0.25 W m–2.
Observations: Oceanic Climate Change and Sea Level Chapter 5
decrease during 1980 to 1983. The 0 to 700 m layer cooled at
a rate of 1.2 W m–2 during this period. Most of this cooling
occurred in the Pacic Ocean and may have been associated
with the reversal in polarity of the PDO (Stephens et al., 2001;
Levitus et al., 2005c, see also Section 3.6.3). Examination of the
geographical distribution of the differences in 0 to 700 m heat
content between the 1977–1981 and 1965–1969 pentads and
the 1986–1990 and 1977–1981 pentads shows that the pattern
of heat content change has spatial scales of entire ocean basins
and is also found in similar analyses by Ishii et al. (2006). The
Pacic Ocean dominates the decadal variations of global heat
content during these two periods. The origin of this variability
is not well understood.
Based on model experiments, it has been suggested that errors
resulting from the highly inhomogeneous distribution of ocean
observations in space and time (see Appendix 5.A.1) could
lead to spurious variability in the analysis (e.g., Gregory et al.,
2004, AchutaRao et al., 2006). As discussed in the appendix,
even in periods with overall good coverage in the observing
system, large regions in Southern Hemisphere (SH) are not well
sampled, and their contribution to global heat content variability
is less certain. However, the large-scale nature of heat content
variability, the similarity of the Levitus et al. (2005a) and the Ishii
et al. (2006) analyses and new results showing a decrease in the
global heat content in a period with much better data coverage
(Lyman et al., 2006), gives condence that there is substantial
inter-decadal variability in global ocean heat content. Implications for Earth’s Heat Balance
To place the changes of ocean heat content in perspective,
Figure 5.4 provides updated estimates of the change in heat
content of various components of the Earth’s climate system for
the period 1961 to 2003 (Levitus et al., 2005a). This includes
changes in heat content of the lithosphere (Beltrami et al., 2002),
the atmosphere (e.g., Trenberth et al., 2001) and the total heat of
fusion due to melting of i) glaciers, ice caps and the Antarctic
and Greenland Ice Sheets (see Chapter 4) and ii) arctic sea ice
(Hilmer and Lemke, 2000). The increase in ocean heat content
is much larger than any other store of energy in the Earth’s heat
balance over the two periods 1961 to 2003 and 1993 to 2003,
and accounts for more than 90% of the possible increase in
heat content of the Earth system during these periods. Ocean
heat content variability is thus a critical variable for detecting
the effects of the observed increase in greenhouse gases in the
Earth’s atmosphere and for resolving the Earth’s overall energy
balance. It is noteworthy that whereas ice melt from glaciers,
ice caps and ice sheets is very important in the sea level budget
Figure 5.3. Linear trend (1955–2003) of zonally averaged temperature in the upper 1,500 m of the water column of the Atlantic, Pacific, Indian and World Oceans. The contour
interval is 0.05°C per decade, and the dark solid line is the zero contour. Red shading indicates values equal to or greater than 0.025°C per decade and blue shading indicates
values equal to or less than –0.025°C per decade. Based on the work of Levitus et al. (2005a).
Chapter 5 Observations: Oceanic Climate Change and Sea Level
(contributing about 40%), the energy associated with ice melt
contributes only about 1% to the Earth’s energy budget.
5.2.3 Ocean Salinity
Ocean salinity changes are an indirect but potentially
sensitive indicator for detecting changes in precipitation,
evaporation, river runoff and ice melt. The patterns of salinity
change can be used to infer changes in the Earth’s hydrological
cycle over the oceans (Wong et al., 1999; Curry et al., 2003) and
are an important complement to atmospheric measurements.
Figure 5.5 shows the linear trends (based on pentadal anomaly
elds) of zonally averaged salinity in the upper 500 m of the
World Ocean and individual ocean basins (Boyer et al., 2005)
from 1955 to 1998. A total of 2.3 million salinity proles were
used in this analysis, about one-third of the amount of data used
in the ocean heat content estimates in Section 5.2.2.
Estimates of changes in the freshwater content of the global
ocean have suggested that the global ocean is freshening (e.g.,
Antonov et al., 2002), however, sampling limitations due to data
sparsity in some regions, particularly the SH, means that such
estimates have an uncertainty that is not possible to quantify.
Between 15°S and 42°N in the Atlantic Ocean there is a
salinity increase in the upper 500 m layer. This region includes
the North Atlantic subtropical gyre. In the 42°N to 72°N region,
including the Labrador, Irminger and Icelandic Seas, there is a
freshening trend (discussed further in Section 5.3). The increase
in salinity north of 72°N (Arctic Ocean) is highly uncertain
because of the paucity of data in this region.
South of 50°S in the polar region of the Southern Ocean,
there is a relatively weak freshening signal. Freshening occurs
throughout most of the Pacic with the exception of the South
Pacic subtropical gyre between 8°S and 32°S and above
300 m where there is an increase in salinity. The near-surface
Indian Ocean is characterised mainly by increasing salinity.
However, in the latitude band 5°S to 42°S (South Indian gyre)
in the depth range of 200 to 1,000 m, there is a freshening of
the water column.
The results shown here document that ocean salinity and
hence freshwater are changing on gyre and basin scales, with the
near-surface waters in the more evaporative regions increasing
in salinity in almost all ocean basins. In the high-latitude
regions in both hemispheres the surface waters are freshening
consistent with these regions having greater precipitation,
although higher runoff, ice melting, advection and changes in
the MOC (Häkkinen, 2002) may also contribute. In addition to
these meridional changes, the Atlantic is becoming saltier over
much of the water column (Figure 5.5 and Boyer et al., 2005).
Although the South Pacic subtropical region is becoming
saltier, on average the whole water column in the Pacic
Basin is becoming fresher (Boyer et al., 2005). The increasing
difference in volume-averaged salinity between the Atlantic
and Pacic Oceans suggests changes in freshwater transport
between these two ocean basins.
We are condent that vertically coherent gyre and basin scale
changes have occurred in the salinity (freshwater content) of
parts of the World Ocean during the past several decades. While
the available data and their analyses are insufcient to identify
in detail the origin of these changes, the patterns are consistent
with a change in the Earth’s hydrological cycle, in particular
with changes in precipitation and inferred larger water transport
in the atmosphere from low latitudes to high latitudes and from
the Atlantic to the Pacic (see Section 3.3.2).
5.2.4 Air-Sea Fluxes and Meridional Transports
The global average changes in ocean heat content discussed
above are driven by changes in the air-sea net energy ux (see
Section At regional scales, few estimates of heat ux
changes have been possible. During the last 50 years, net heat
uxes from the ocean to the atmosphere demonstrate locally
decreasing values (up to 1 W m–2 yr–1) over the southern ank
of the Gulf Stream and positive trends (up to 0.5 W m–2 yr–1) in
the Atlantic central subpolar regions (Gulev et al., 2006). At the
global scale, the accuracy of the ux observations is insufcient
to permit a direct assessment of changes in heat ux. Air-sea
uxes are discussed in Section 3.5.6.
Figure 5.4. Energy content changes in different components of the Earth system
for two periods (1961–2003 and 1993–2003). Blue bars are for 1961 to 2003,
burgundy bars for 1993 to 2003. The ocean heat content change is from this section
and Levitus et al. (2005c); glaciers, ice caps and Greenland and Antarctic Ice Sheets
from Chapter 4; continental heat content from Beltrami et al. (2002); atmospheric en-
ergy content based on Trenberth et al. (2001); and arctic sea ice release from Hilmer
and Lemke (2000). Positive energy content change means an increase in stored
energy (i.e., heat content in oceans, latent heat from reduced ice or sea ice volumes,
heat content in the continents excluding latent heat from permafrost changes, and
latent and sensible heat and potential and kinetic energy in the atmosphere). All error
estimates are 90% confidence intervals. No estimate of confidence is available for
the continental heat gain. Some of the results have been scaled from published re-
sults for the two respective periods. Ocean heat content change for the period 1961
to 2003 is for the 0 to 3,000 m layer. The period 1993 to 2003 is for the 0 to 700 m
(or 750 m) layer and is computed as an average of the trends from Ishii et al. (2006),
Levitus et al. (2005a) and Willis et al. (2004).
Observations: Oceanic Climate Change and Sea Level Chapter 5
Estimates of the climatological mean oceanic meridional heat
transport derived from atmospheric observations (e.g., Trenberth
and Caron, 2001) and from oceanographic cross sections (e.g.,
Ganachaud and Wunsch, 2003) are in fair agreement, despite
considerable uncertainties (see Appendix 5.A.2). The ocean heat
transport estimate derived from integration of climatological
air-sea heat ux elds (e.g., Grist and Josey, 2003) is in good
agreement with an independent oceanographic cross section
at 32°S. Estimates of changes in the Atlantic meridional heat
transport are discussed in Section 5.3.2.
5.3 Regional Changes in Ocean
Circulation and Water Masses
5.3.1 Introduction
Robust long-term trends in global- and basin-scale ocean heat
content and basin-scale salinity were shown in Section 5.2. The
observed heat and salinity trends are linked to changes in ocean
circulation and other manifestations of global change such as
oxygen and carbon system parameters (see Section 5.4). Global
ocean changes result from regional changes in these properties,
assessed in this section. Evidence for change in temperature,
salinity and circulation is described globally and then for each
of the major oceans. Two marginal seas with multi-decadal time
series are also examined as examples of regional variations.
The upper ocean in all regions is close to the atmospheric
forcing and has the largest variability; it is also the best sampled.
For these reasons, Section 5.2 mainly assessed upper-ocean
observations for long-term trends in heat content and salinity.
However, there are important changes in heat and salinity at
intermediate and abyssal depths, restricted to regions that are
relatively close to the main sources of deep and intermediate
waters. These sources are most vigorous in the northern North
Atlantic and the Southern Ocean around Antarctica. This is
illustrated well in salinity differences shown for the Atlantic
(1985–1999 minus 1955–1969) and Pacic (1980s minus
1960s) in Figure 5.6. Striking changes in salinity are found
from the surface to the bottom in the northern North Atlantic
near water mass formation sites that ll the water column
(Section 5.3.2); bottom changes elsewhere are small, being most
prevalent at the under-sampled southern ends of both sections.
At mid-depth (500 to 2,000 m), the Atlantic and southern end
of the Pacic section show widespread change, but the North
Pacic signal is weaker and shallower because it has only weak
intermediate water formation (and no deep water formation).
Changes in intermediate and deep waters can ultimately affect
Figure 5.5: Linear trends (1955–1998) of zonally averaged salinity (psu) in the upper 500 m of the Atlantic, Pacific, Indian and World Oceans. The contour interval is 0.01 psu
per decade and dashed contours are ±0.005 psu per decade. The dark solid line is the zero contour. Red shading indicates values equal to or greater than 0.005 psu per decade
and blue shading indicates values equal to or less than –0.005 psu per decade. Based on the work of Boyer et al. (2005).
Chapter 5 Observations: Oceanic Climate Change and Sea Level
the ocean’s vertical stratication
and overturning circulation; the
topic of the overturning circulation
in the North Atlantic is considered
in Section 5.3.2.
The observed changes in salinity
are of global scale, with similar
patterns in different ocean basins
(Figure 5.6). The subtropical waters
have increased in salinity and the
subpolar surface and intermediate
waters have freshened in both the
Atlantic and Pacic Oceans during
the period from the 1960s to the
1990s and in both hemispheres in
each ocean. The waters that underlie
the near-surface subtropical waters
have freshened due to equatorward
circulation of the freshened subpolar
surface waters; in particular, the fresh
intermediate water layer (at ~1,000
m) in the SH has freshened in both
the Atlantic and Pacic Oceans. In
the Northern Hemisphere (NH), the
Pacic intermediate waters have
freshened, and the underlying deep
waters did not change, consistent
with no local bottom water source
in the North Pacic. In the central
North Atlantic, the intermediate
layer (approximately 900–1,200 m)
became saltier due to increased
salinity in the outow from the
Mediterranean that feeds this layer.
5.3.2 Atlantic and Arctic
The North Atlantic Ocean has
a special role in long-term climate
assessment because it is central to one of the two global-scale
MOCs (see Box 5.1), the other location being the Southern
Ocean. The long-term trends in depth-integrated Atlantic heat
content for the period 1955 to 2003 (Figure 5.2) are broadly
consistent with the warming tendencies identied from the
global analyses of SST (see Section The subtropical
gyre warmed and the subpolar gyre cooled over that period,
consistent with a predominantly positive phase of the NAO
during the last several decades. The warming extended down to
below 1,000 m, deeper than anywhere else in the World Ocean
(Figure 5.3 Atlantic), and was particularly pronounced under the
Gulf Stream and North Atlantic Current near 40°N. Long-term
trends in salinity towards freshening in the subpolar regions
and increased salinity in the subtropics through the mid-1990s
(Figure 5.5 Atlantic and Figure 5.6a) are consistent with the
global tendencies for freshening of relatively fresher regions
and increased salinity in saltier regions (Section 5.2.3).
Figure 5.6. Meridional sections of differences in salinity (psu) of the a) Atlantic
Ocean for the period 1985 to 1999 minus 1955 to 1969 and b) Pacific Ocean for
the World Ocean Circulation Experiment (WOCE) 150°W section (1991–1992) and
historical data from 1968 plus or minus 7.5 years. Contours are the mean salinity
fields along each section and show the key features. The salinity differences are dif-
ferences along isopycnals that have been mapped to pressure surfaces. The Atlantic
section is along the western side of the Atlantic Ocean and the Pacific section is
along 150°W. The two figures are redrafted from Curry et al. (2003) and Wong et al
(2001). Water masses shown include Antarctic Intermediate Water (AAIW), Circum-
polar Deep Water (CDW), North Atlantic Deep Water (NADW), Mediterranean Water
(MW), Labrador Sea Water (LSW), Denmark Strait Overflow Water (DSOW) and North
Pacific Intermediate Water NPIW). The areas shaded in grey represent the seafloor
and oceanic crust.
Observations: Oceanic Climate Change and Sea Level Chapter 5
Figure 5.7. The longest available time series of salinity (psu; upper panel) and potential temperature (°C, lower
panel) in the central Labrador Sea from 1949 to 2005 (updated from Yashayaev et al., 2003). The dashed lines are
contours of potential density (kg m-3, difference from 1,000 kg m-3) and are the same on both panels. North Atlantic Subpolar Gyre, Labrador Sea and
Nordic Seas
In the North Atlantic subpolar gyre, Labrador Sea and
Nordic Seas, large salinity changes have been observed that
have been associated with changed inputs of fresh water (ice
melt, ocean circulation and river runoff) and with the NAO.
Advection of these surface and deep salinity anomalies has
been traced around the whole subpolar gyre including the
Labrador and Nordic Seas. These anomalies are often called
‘Great Salinity Anomalies’ (GSAs; e.g., Dickson et al., 1988;
Belkin, 2004). During a positive phase of the NAO, the subpolar
gyre strengthens and expands towards the east, resulting in
lower surface salinity in the central subpolar region (Levitus,
1989; Reverdin et al., 1997; Bersch, 2002). Three GSAs have
been thoroughly documented: one from 1968 to 1978, one in
the 1980s and one in the 1990s. Observational and modelling
studies show that the relative inuence of local events and
advection differ between different GSA events and regions
(Houghton and Visbeck, 2002; Josey and Marsh, 2005).
These surface salinity anomalies have affected the Labrador
Sea and the production of Labrador Sea Water (LSW), a major
component of the North Atlantic Deep Water (NADW) and
contributor to the lower limb of the MOC. The LSW appears
to alternate between dense, cold types and less dense, warm
types (Yashayaev et al., 2003; Kieke et al., 2006) possibly with
more production of dense LSW during years of positive-phase
NAO (Dickson et al., 1996). Since 1965 to 1970, the LSW
has had a signicant freshening trend with a superimposed
variability consisting of three saltier periods, coinciding with
warmer water, and two freshening and cooling periods in the
1970s and 1990s (Figure 5.7). During the period 1988 to 1994,
an exceptionally large volume of
cold, fresh and dense LSW was
produced (Sy et al., 1997; Lazier
et al., 2002), unprecedented in the
sparse time series that extends back
to the 1930s (Talley and McCartney,
1982). The Labrador Sea has now
returned to a warmer, more saline
state; most of the excess volume
of the dense LSW has disappeared,
the mid-layers became warmer and
saltier, and the production of LSW
shifted to the warmer type (e.g.,
Lazier et al., 2002; Yashayaev et al.,
2003; Stramma et al., 2004). This
warming and increased salinity and
reduction in LSW was associated
with the weakening of the North
Atlantic subpolar gyre, seen also in
satellite altimetry data (Häkkinen
and Rhines, 2004).
The eastern half of the subpolar
North Atlantic also freshened through
the 1980s and into the 1990s, but the
upper ocean has been increasing in
salinity or remaining steady since
then, depending on location. About
two-thirds of the freshening in this
region has been attributed to an
increase in precipitation associated
with a climate pattern known as
the East Atlantic Pattern (Josey and
Marsh, 2005), with the NAO playing
a secondary role. From 1965 to 1995,
the subpolar freshening amounted
to an equivalent freshwater layer of
approximately 3 m spread evenly
over its total area (Belkin, 2004;
Curry and Mauritzen, 2005).
Chapter 5 Observations: Oceanic Climate Change and Sea Level
Subsurface salinity in the Nordic Seas has also decreased
markedly since the 1970s (Dickson et al., 2003), directly
affecting the salinity of the Nordic Sea overow waters that
contribute to NADW. This decrease in subsurface salinity was
associated with lower salinity of the Atlantic waters entering
the Nordic seas and related to the high NAO index and
intensication of the subpolar gyre. Since 1994, the salinity of
the inow from the North Atlantic has been increasing, reaching
the highest values since 1948, largely due to a weakening of
the subpolar gyre circulation that allowed more warm water
into the Nordic Seas, associated with a decreasing NAO index
(Hátún et al., 2005).
The densest waters contributing to NADW and to the deep
limb of the MOC arise as overows from the upper 1,500 m of
the Nordic Seas through the Denmark Strait and Faroe Channel.
The marked freshening of the overow water masses exiting
the Arctic was associated with growing sea ice export from
the Arctic and precipitation in the Nordic Seas (Dickson et al.,
2002, 2003). The transports of the overow waters, of which the
largest component is through Denmark Strait, have varied by
about 30% (Macrander et al., 2005), but there has been no clear
trend in this location. Overall, the overows that contribute to
NADW from the Nordic Seas have remained constant to within
the known variability.
The overall pattern of change in the North Atlantic subpolar
gyre is one of a trend towards fresher values over most of the
water column from the mid-1960s until the mid-1990s. Since
then, there has been a return to warmer and more saline waters
Box 5.1: Has the Meridional Overturning Circulation in the Atlantic Changed?
The global Meridional Overturning Circulation consists primarily of dense waters that sink to the abyssal ocean at high latitudes
in the North Atlantic Ocean and near Antarctica. These dense waters then spread across the equator with comparable ows of ap-
proximately 17 and 14 Sv (106 m3 s–1), respectively (Orsi et al., 2002; Talley et al., 2003a). The North Atlantic overturning circulation
(henceforth ‘MOC’) is characterised by an inow of warm, saline upper-ocean waters from the south that gradually increase in density
from cooling as they move northward through the subtropical and subpolar gyres. They also freshen, which reduces the density in-
crease. The inows reach the Nordic Seas (Greenland, Iceland and Norwegian Seas) and the Labrador Sea, where they are subject to
deep convection, sill overows and vigorous mixing. Through these processes NADW is formed, constituting the southward-owing
lower limb of the MOC.
Climate models show that the Earth’s climate system responds to changes in the MOC (e.g., Vellinga and Wood, 2002), and also
suggest that the MOC might gradually decrease in transport in the 21st century as a consequence of anthropogenic warming and ad-
ditional freshening in the North Atlantic (Bi et al., 2001; Gregory et al., 2005; see also Chapter 10). However, observations of changes in
the MOC strength and variability are fragmentary; the best evidence for observational change comes from the North Atlantic.
There is evidence for a link between the MOC and abrupt changes in surface climate during the past 120 kyr, although the exact
mechanism is not clear (Clark et al., 2002). At the end of the last glacial period, as the climate warmed and ice sheets melted, there
were a number of abrupt oscillations, for example, the Younger Dryas and the 8.2 ka cold event (see Section 6.4), which may have been
caused by changes in ocean circulation. The variability of the MOC during the Holocene after the 8.2 ka cooling event is clearly much
smaller than during glacial times (Keigwin et al., 1994; see Section 6.4).
Observed changes in MOC transport, water properties and water mass formation are inconclusive about changes in the MOC
strength (see Section This is partially due to decadal variability and partially due to inadequate long-term observations. From
repeated hydrographic sections in the subtropics, Bryden et al. (2005) concluded that the MOC transport at 25°N had decreased by
30% between 1957 and 2004, but the presence of signicant unsampled variability in time and the lack of supporting direct current
measurements reduces condence in this estimate. Direct measurements of the two major sill overows have shown considerable
variability in the dominant Denmark Strait Overow without enough years of coverage to discern long-term trends (Macrander et al.,
2005). The observed freshening of the overows and the associated reduction in density from 1965 to 2000 (see Section 5.3.2) has so
far not led to a signicant weakening of the MOC (Dickson et al., 2003; Curry and Mauritzen, 2005). Moreover, large decadal variability
observed since 1960 in salinity and temperature of the surface waters, including the recent increase in salinity of the surface waters
feeding the MOC, obscures the long-term trend (Hátún et al., 2005; ICES 2005) and hence conclusions about potential MOC changes.
Changes in the MOC can also be caused by changes in Labrador Sea convection, with strong convection corresponding to higher
MOC. Convection was strong from the 1970s to 1995, but thereafter the Labrador Sea warmed and re-stratied (Lazier et al., 2002;
Yashayaev et al., 2003) and convection has been weaker. Based on observed SST patterns, it was concluded that the MOC transport has
increased by about 10% from 1970 to the 1990s (Knight et al., 2005; Latif et al., 2006). From direct current meter observations at the exit
of the subpolar North Atlantic, Schott et al. (2004) concluded that the Deep Water outow, while varying at shorter time scales, had no
signicant trend during the 1993 to 2001 period.
In summary, it is very likely that up to the end of the 20th century the MOC was changing signicantly at interannual to decadal
time scales. Given the above evidence from components of the MOC as well as uncertainties in the observational records, over the
modern instrumental record no coherent evidence for a trend in the mean strength of the MOC has been found.
Observations: Oceanic Climate Change and Sea Level Chapter 5
(Figure 5.7), which coincides with the change in NAO and East
Atlantic Pattern. However, this return to saltier waters has not
been sustained for a long enough period to change the sign of
the long-term trends (Figure 5.5 Atlantic). Arctic Ocean
Climate change in the Arctic Ocean and Nordic Seas is
closely linked to the North Atlantic subpolar gyre (Østerhus et
al., 2005). Within the Arctic Ocean and Nordic Seas, surface
temperature has increased since the mid-1980s and continues
to increase (Comiso, 2003). In the Atlantic waters entering
the Nordic Seas, a temperature increase in the late 1980s and
early 1990s (Quadfasel et al., 1991; Carmack et al., 1995) has
been associated with the transition in the 1980s towards more
positive NAO states. Warm Atlantic waters have also been
observed to enter the Arctic as pulses via Fram Strait and then
along the slope to the Laptev Sea (Polyakov et al., 2005); the
increased heat content and increased transport in the pulses
both contribute to net warming of the arctic waters (Schauer
et al., 2004). Multi-decadal variability in the temperature of
the Atlantic Water core affecting the top 400 m in the Arctic
Ocean has been documented (Polyakov et al., 2004). Within the
Arctic, salinity increased in the upper layers of the Amundsen
and Makarov Basins, while salinity of the upper layers in the
Canada Basin decreased (Morison et al., 1998). Compared to the
1980s, the area of upper waters of Pacic origin has decreased
(McLaughlin et al., 1996; Steele and Boyd, 1998).
During the 1990s, changed winds caused eastward redirection
of river runoff from the Laptev Sea (Lena River, etc.), reducing
the low-salinity surface layer in the central Arctic Ocean (Steele
and Boyd, 1998), thus allowing greater convection and heat
transport into the surface arctic layer from the more saline
subsurface Atlantic layer. Thereafter, however, the stratication
in the central Arctic (Amundsen Basin) increased and a low-
salinity mixed layer was again observed at the North Pole in
2001, possibly due to a circulation change that restored the
river water input (Björk et al., 2002). Circulation variability that
shifts the balance of fresh and saline surface waters in the Arctic,
with associated changes in sea ice, might be associated with the
NAM (Proshutinsky and Johnson, 1997; Rigor et al., 2002),
however, the long-term decline in arctic sea ice cover appears
to be independent of the NAM (Comiso, 2002). While there is
signicant decadal variability in the Arctic Ocean, no systematic
long-term trend in subsurface arctic waters has been identied. Subtropical and Equatorial Atlantic
In the North Atlantic subtropical gyre, circulation, SST, the
thickness of near-surface Subtropical Mode Water (STMW,
Hanawa and Talley, 2001) and thermocline ventilation are all
highly correlated with the NAO, with some time lags. A more
positive NAO state, with westerlies shifted northwards, results
in a decreased Florida Current transport (Baringer and Larsen,
2001), a likely delayed northward shift of the Gulf Stream
position (Joyce et al., 2000; Seager et al., 2001; Molinari,
2004), and decreased subtropical eddy variability (Penduff et
al., 2004). In the STMW, low thickness and production and
higher temperature result from a high NAO index (e.g., Talley,
1996; e.g., Hazeleger and Drijfhout, 1998; Marsh, 2000). The
volume of STMW is likely to lag changes in the NAO by two
to three years, and low (high) volumes are associated with high
(low) surface layer temperatures because of changes in both
convective forcing and location of STMW formation. While
quasi-cyclic variability in STMW renewal is apparent over the
1960 to 1980 period, the total volume of STMW has remained
low through 2000 since a peak in 1983 to 1984, associated with
a relatively persistent positive NAO phase during the late 1980s
and early 1990s (Lazier et al., 2002; Kwon and Riser, 2004).
In the subtropics at depths of 1,000 to 2,000 m, the
temperature has increased since the late 1950s at Bermuda, at
24°N, and at 52°W and 66°W in the Gulf Stream (Bryden et
al., 1996; Joyce and Robbins, 1996; Joyce et al., 1999). These
warming trends reect reduced production of LSW (Lazier,
1995) and increased salinity and temperature of the waters
from the Mediterranean (Roether et al., 1996; Potter and Lozier,
2004). After the mid-1990s at greater depths (1,500–2,500 m),
temperature and salinity decreased, reversing the previous
warming trend, most likely due to delayed appearance of the
new colder and fresher Labrador Sea Water produced in the
Intermediate water (800–1,200 m) in the mid-latitude
eastern North Atlantic is strongly inuenced by the saline
Mediterranean Water (MW; Section This saline layer
joins the southward-owing NADW and becomes part of it in
the tropical Atlantic. This layer has warmed and become more
saline since at least 1957 (Bryden et al., 1996), continuing
during the last decade (1994–2003) at a rate of more than
0.2°C per decade with a rate of 0.4°C per decade at some
levels (Vargas-Yáñez et al., 2004). In the Bay of Biscay (44°N;
González-Pola et al., 2005) and at Gibraltar (Millot et al.,
2006), similar warming was observed through the thermocline
and into the core of the MW. From 1955 to 1993, the trend was
about 0.1°C per decade in a zone west of Gibraltar (Potter and
Lozier, 2004), and of almost the same magnitude even west of
the mid-Atlantic Ridge (Curry et al., 2003).
Surface waters in the Southern Ocean, including the high-
latitude South Atlantic, set the initial conditions for bottom
water in the (SH). This extremely dense Antarctic Bottom
Water (AABW), which is formed around the coast of Antarctica
(see Section, spreads equatorward and enters the Brazil
Basin through the narrow Vema Channel of the Rio Grande
Rise at 31°S. Ongoing observations of the lowest bottom
temperatures there have revealed a slow but consistent increase
of the order 0.002°C yr–1 in the abyssal layer over the last 30
years (Hogg and Zenk, 1997).
In the tropical Atlantic, the surface water changes are partly
associated with the variability of the marine Inter-tropical
Convergence Zone, which has strong seasonal variability
(Mitchell and Wallace, 1992; Biasutti et al., 2003; Stramma et
al., 2003). Tropical Atlantic variability on interannual to decadal
time scales can be inuenced by a South Atlantic dipole in SST
Chapter 5 Observations: Oceanic Climate Change and Sea Level
(Venegas et al., 1998), associated with latent heat uxes related
to changes in the subtropical high (Sterl and Hazeleger, 2003).
The South Equatorial Current provides a region for subduction
of the water masses (Hazeleger et al., 2003) and may also
maintain a propagation pathway for water mass anomalies
towards the north (Lazar et al., 2002).
The North Atlantic Oscillation is an important driver of
the oceanic water mass variations in the upper North Atlantic
subtropical gyre. Its effects are also observed at depths greater
than 1,500 m within the subtropical gyre consistent with the
large-scale circulation and changes in source waters in the North
Atlantic Ocean. While there are coherent changes in the long-
term trends in temperature and salinity (Section 5.2), decadal
variations are an important climate signal for this region. Mediterranean Sea
Marked changes in thermohaline properties have been
observed throughout the Mediterranean (Manca et al., 2002).
In the western basin, the Western Mediterranean Deep Water
(WMDW), formed in the Gulf of Lions, warmed during the
last 50 years, interrupted by a short period of cooling in the
early 1980s, the latter reected in cooling of the Levantine
Intermediate Water between the late 1970s and mid-1980s
(Brankart and Pinardi, 2001). The WMDW warming is in
agreement with recent atmospheric temperature changes over
the Mediterranean (Luterbacher et al., 2004). The salt content
of the WMDW has also been steadily increasing during the
last 50 years, mainly attributed to decreasing precipitation
over the region since the 1940s (Krahmann and Schott, 1998;
Mariotti et al., 2002) and to anthropogenic reduction in the
freshwater inow (Rohling and Bryden, 1992). These changes
in water properties and circulation are linked to the long-term
variability of surface uxes (Krahmann and Schott, 1998)
with contributions from the NAO (Vignudelli et al., 1999) that
produce consistent changes in surface heat uxes and a net
warming of the Mediterranean Sea (Rixen et al., 2005).
These changes in the temperature and salinity within the
Mediterranean have affected the outow of water into the
North Atlantic at Gibraltar (see also Section Part of
this shift in Mediterranean outow properties has been traced
to the Eastern Mediterranean. During 1987 to 1991, the Eastern
Mediterranean Deep Water became warmer and saltier due to
the switch of its source water from the Adriatic to the Aegean
(Klein et al., 2000; Gertman et al., 2006), most likely related to
changes in the heat and freshwater ux anomalies in the Aegean
Sea (Tsimplis and Rixen, 2002; Josey, 2003; Rupolo et al.,
2003). This 1987 to 1991 switch of source waters has continued
and increased its impact, with density of the westward outow
in Sicily Strait now denser (Gasparini et al., 2005). While there
are strong natural variations in the Mediterranean, overall
there is a discernible trend of increased salinity and warmer
temperature in key water masses over the last 50 years and this
signal is observable in the North Atlantic.
5.3.3 Pacific Ocean
The upper Pacic Ocean has been warming and freshening
overall, as revealed in global heat and freshwater analyses
(Section 5.2, Figure 5.5). The subtropical North and South
Pacic have been warming. In the SH, the major warming
footprint is associated with the thick mode waters north of the
Antarctic Circumpolar Current. The North Pacic has cooled
along 40°N. Long-term trends are rather difcult to discern
in the upper Pacic Ocean because of the strong interannual
and decadal variability (ENSO and the PDO) and the relatively
short length of the observational records. Changes associated
with ENSO are described in Section 3.6.2 and are not included
here. Overall, the Pacic is freshening but there are embedded
salinity increases in the subtropical upper ocean, where strong
evaporation dominates. PacicUpperOceanChanges
In the North Pacic, the zonally averaged temperature
warming trend from 1955 to 2003 (Figure 5.3) is dominated by
the PDO increase in the mid-1970s. The strong cooling between
50 and 200 m is due to relaxation and subsequent shallowing of
the tropical thermocline, resulting from a decrease in the shallow
tropical MOC and a relaxation of the equatorial thermocline
(McPhaden and Zhang, 2002), although after 1998 this shallow
overturning circulation returned to levels almost as high as in
the 1970s (McPhaden and Zhang, 2004).
Warming in the North Pacic subtropics, cooling around
40°N and slight warming farther north is the pattern associated
with a positive PDO (strengthened Aleutian Low; Miller and
Douglas, 2004; see Figure 3.28). Within the North Pacic Ocean,
a positive PDO state such as occurred after 1976 is characterised
by a strengthened Kuroshio Extension. After 1976, the Kuroshio
Extension and North Pacic Current transport increased by 8%
and expanded southward (Parrish et al., 2000). The Kuroshio’s
advection of temperature anomalies has been shown to be of
similar importance to variations in ENSO and the strength of
the Aleutian Low in maintaining the positive PDO (Schneider
and Cornuelle, 2005). The Oyashio penetrated farther southward
along the coast of Japan during the 1980s than during the
preceding two decades, consistent with a stronger Aleutian Low
(Sekine, 1988; Hanawa, 1995; Sekine, 1999). A shoaling of
the halocline in the centre of the western subarctic gyre and a
concurrent southward shift of the Oyashio extension front during
1976 to 1998 vs. 1945 to 1975 has been detected (Joyce and
Dunworth-Baker, 2003). Similarly, mixed layer depth decreased
throughout the eastern subarctic gyre, with a distinct trend over
50 years (Freeland et al., 1997; Li et al., 2005).
Temperature changes in upper-ocean water masses in
response to the more positive phase of the PDO after 1976
are well documented. The thick water mass just south of the
Kuroshio Extension in the subtropical gyre (Subtropical Mode
Water) warmed by 0.8°C from the mid-1970s to the late 1980s,
associated with stronger Kuroshio advection, and the thick
water mass along the subtropical-subpolar boundary near 40°N
Observations: Oceanic Climate Change and Sea Level Chapter 5
(North Pacic Central Mode Water) cooled by 1°C following
the shift in the PDO after 1976 (Yasuda et al., 2000; Hanawa
and Kamada, 2001).
Trends towards increased heat content include a major
signal in the subtropical South Pacic, within the thick mixed
layers just north of the Antarctic Circumpolar Current (Willis
et al., 2004; Section 5.3.5). The strength of the South Pacic
subtropical gyre circulation increased more than 20% after
1993, peaking in 2003, and subsequently declined. This spin up
is linked to an increase of Ekman pumping over the gyre due to
an increase in the SAM index (Roemmich et al., 2007).
The marginal seas of the Pacic Ocean are also subject to
climate variability and change. Like the Mediterranean in the
North Atlantic, the Japan (or East) Sea is nearly completely
isolated from the adjacent ocean basin, and forms all of its
own waters beneath the shallow pycnocline. Because of this
sea’s limited size, it responds quickly through its entire depth
to surface forcing changes. The warming evident through the
global ocean is clearly apparent in this isolated basin, which
warmed by 0.1°C at 1,000 m and 0.05°C below 2,500 m since
the 1960s. Salinity at these depths also changed, by 0.06 psu
per century for depths of 300 to 1,000 m and by –0.02 psu per
century below 1,500 m (Kwon et al., 2004). These changes
have been attributed to reduced surface heat loss and increased
surface salinity, which have changed the mode of ventilation
(Kim et al., 2004). Deep water production in the Japan (East)
Sea slowed for many decades, with a marked decrease in
dissolved oxygen from the 1930s to 2000 at a rate of about
0.8 µmol kg–1 yr–1 (Gamo et al., 1986; Minami et al., 1998).
However, possibly because of weakened vertical stratication
at mid-depths associated with the decades-long warming, deep-
water production reappeared after the 2000–2001 severe winter
(e.g., Kim et al., 2002; Senjyu et al., 2002; Talley et al., 2003b).
Nevertheless, the overall trend has continued with lower deep-
water production in subsequent years. Intermediate and Deep Circulation and Water
Property Changes
Since the 1970s, the major mid-depth water mass in the
North Pacic, North Pacic Intermediate Water (NPIW), has
been freshening and has become less ventilated, as measured by
oxygen content (see Section 5.4.3). The NPIW is formed in the
subpolar North Pacic, with most inuence from the Okhotsk
Sea, and reects changes in northern North Pacic surface
conditions. The salinity of NPIW decreased by 0.1 and 0.02
psu in the subpolar and subtropical gyres, respectively (Wong
et al., 2001; Joyce and Dunworth-Baker, 2003). An oxygen
decrease and nutrient increase in the NPIW south of Hokkaido
from 1970 to 1999 was reported (Ono et al., 2001), along with
a subpolar basin-wide oxygen decrease from the mid-1980s to
the late 1990s (Watanabe et al., 2001). Warming and freshening
occurred in the Okhotsk Sea in the latter half of the 20th
century (Hill et al., 2003). The Okhotsk Sea intermediate water
thickness was reduced and its density decreased in the 1990s
(Yasuda et al., 2001).
In the southwest Pacic, in the deepest waters originating
from the North Atlantic and Antarctica, cooling and freshening
of 0.07°C and 0.01 psu from 1968 to 1991 was observed
(Johnson and Orsi, 1997) and attributed to a change in the
relative importance of Antarctic and North Atlantic source
waters and weakening bottom transport. Bottom waters in the
North Pacic are farther from the surface sources than any other
of the world’s deep waters. They are also the most uniform, in
terms of spatial temperature and salinity variations. A large-
scale, signicant warming of the bottom 1,000 m across the
entire North Pacic of the order of 0.002°C occurred between
1985 and 1999, measurable because of the high accuracy of
modern instruments (Fukasawa et al., 2004). The cause of this
warming is uncertain, but could have resulted from warming of
the deep waters in the South Pacic and Southern Ocean, where
mid-depth changes since the 1950s are as high as 0.17°C (Gille,
2002; see Figure 5.8), and/or from the declining bottom water
transport into the deep North Pacic (Johnson et al., 1994).
5.3.4 Indian Ocean
The upper Indian Ocean has been warming everywhere
except in a band centred at about 12°S (South Equatorial
Current), as seen in Section 5.2 (Figure 5.3). In the tropical and
eastern subtropical Indian Ocean (north of 10°S), warming in the
upper 100 m (Qian et al., 2003) is consistent with the signicant
warming of the sea surface from 1900 to 1999 (see Section 3.2.2
and Figure 3.9). The surface warming trend during the period
1900 to 1970 was relatively weak, but increased signicantly
in the 1970 to 1999 period, with some regions exceeding 0.2°C
per decade.
The global-scale circulation includes transport of warm,
relatively fresh waters from the Pacic passing through the
Indonesian Seas to the Indian Ocean and then onward into
the South Atlantic. Much of this throughow occurs in the
tropics south of the equator, and is strongly affected by ENSO
and the Indian Ocean Dipole (see Section The latter
causes pronounced thermocline variability (Qian et al., 2003)
and includes propagation of upper-layer thickness anomalies
by Rossby waves (Xie et al., 2002; Feng and Meyers, 2003;
Yamagata et al., 2004) in the 3°S to 15°S latitude band that
includes the westward-owing throughow water.
Long-term trends in transport and properties of the
throughow have not been reported. The mean transport into
the Indonesian Seas measured at Makassar Strait from 1996 to
1998 was 9 to 10 Sv (Vranes et al., 2002), matching transports
exiting the Indonesian Seas (e.g., Sprintall et al., 2004). Large
variability in this transport is associated with varying tropical
Pacic and Indian winds (Wijffels and Meyers, 2004), including
a strong ENSO response (e.g., Meyers, 1996), and may be
associated with changes in SST in the tropical Indian Ocean.
Models suggest that upper-ocean warming in the south
Indian Ocean can be attributed to a reduction in the southeast
trade winds and associated decrease in the southward transport
of heat from the tropics to the subtropics (Lee, 2004). The
export of heat from the northern Indian Ocean to the south
Chapter 5 Observations: Oceanic Climate Change and Sea Level
across the equator is accomplished by a wind-driven, shallow
cross-equatorial cell; data assimilation analysis has shown a
signicant decadal reduction in the mass exchange during 1950
to 1990 but little change in heat transport (Schoenefeldt and
Schott, 2006).
Changes in Indian subtropical gyre circulation since the
1960s include a 20% slowdown from 1962 to 1987 (Bindoff
and McDougall, 2000) and a 20% speedup from 1987 to
2002 (Bryden et al., 2003; McDonagh et al., 2005), with the
speedup mainly between 1995 and 2002 (Palmer et al., 2004).
The upper thermocline warmed during the slowdown, and
then cooled during speedup. Simulations of this region and the
analysis of climate change scenarios show that the slowdown
and speedup were part of an oscillatory pattern in the upper
part of this gyre over periods of decades (Murray et al., 2007;
Stark et al., 2006). On the other hand, the lower thermocline
(<10°C) freshened and warmed from 1936 to 2002 (Bryden et
al., 2003), consistent with heat content increases discussed in
Section 5.2 and earlier results.
5.3.5 Southern Ocean
The Southern Ocean, which is the region south of 30°S,
connects the Atlantic, Indian and Pacic Oceans together,
allowing inter-ocean exchange. This region is active in the
formation and subduction of waters that contributed strongly
to the storage of anthropogenic carbon and heat (see Section
5.2). It is also the location of the densest part of the global
overturning circulation, through formation of bottom waters
around Antarctica, fed by deep waters from all of the oceans
to the north. Note that some observed changes found in the
Atlantic, Indian and Pacic Oceans are related to changes in
the Southern Ocean waters but have largely been described in
those sections. Upper-OceanPropertyChanges
The upper ocean in the SH has warmed since the 1960s,
dominated by changes in the thick near-surface layers called
Subantarctic Mode Water (SAMW), located just north of the
Antarctic Circumpolar Current (ACC) that encircles Antarctica.
The observed warming of SAMW is consistent with the
subduction of warmer surface waters from south of the ACC
(Wong et al., 2001; Aoki et al., 2003). In the Upper Circumpolar
Deep Water (UCDW) in the Indian and Pacic sectors of the
Southern Ocean, temperature and salinity have been increasing
(on density surfaces) and oxygen has been decreasing between
the Subantarctic Front near 45°S and the Antarctic Divergence
near 60°S (Aoki et al., 2005a). These changes just below the
mixed layer (~100 to 300 m) are consistent with the mixing of
warmer and fresher surface waters with UCDW, suggesting an
increase in stratication in the surface layer of this polar region.
Mid-depth waters of the Southern Ocean have also warmed in
recent decades. As shown in Figure 5.8, temperatures increased
near 900 m depth between the 1950s and the 1980s throughout
most of the Southern Ocean (Aoki et al., 2003; Gille, 2004).
The largest changes are found near the Antarctic Circumpolar
Current, where the warming at 900 m depth is similar in
magnitude to the increase in regional surface air temperatures.
Analysis of altimeter and Argo oat prole data suggests that,
over the last 10 years, the zonally averaged warming in the
upper 400 m of the ocean near 40°S (Willis et al. 2004) is much
larger than that seen in long-term trends (see Section 5.2, Figure
5.3 World). The warming results from these analyses have been
attributed to a southward shift and increased intensity of the SH
westerlies, which would shift the ACC slightly southward and
intensify the subtropical gyres (e.g., Cai, 2006).
The major mid-depth water mass in the SH, Antarctic
Intermediate Water (AAIW), has also been freshening since the
1960s (Wong et al., 1999; Bindoff and McDougall, 2000; see
Figure 5.6). The Atlantic freshening of AAIW is also supported
by direct observations of a freshening of southern surface waters
(Curry et al., 2003). Antarctic Regions and Antarctic Circumpolar
The ACC, the longest current system in the world, has
a transport through Drake Passage of about 130 Sv, with
signicant interannual variability. Measurements over 25 years
across Drake Passage show no evidence for a systematic trend
in total volume transport between the 1970s and the present
(Cunningham et al., 2003), although continuous subsurface
pressure measurements suggest that trends in seasonality of
transport are highly correlated with similar trends in the SAM
index (Meredith and King, 2005).
There is growing evidence for the changes in the AABW and
intermediate depth waters around Antarctica. In the Weddell
Sea, the deep and bottom water properties varied in the 1990s
Figure 5.8. Temperature trends (°C yr–1) at 900 m depth using data collected from
the 1930s to 2000, including shipboard profile and Autonomous LAgrangian Current
Explorer float data. The largest warming occurs in subantarctic regions, and a slight
cooling occurs to the north. From Gille (2002).
Observations: Oceanic Climate Change and Sea Level Chapter 5
(Robertson et al., 2002; Fahrbach et al., 2004). Changes in
bottom water properties have also been observed downstream
of these source regions (Hogg, 2001; Andrie et al., 2003) and in
the South Atlantic (Section The upper ocean adjacent
to the West Antarctic Peninsula warmed by more than 1°C and
became more saline by 0.25 psu from 1951 to 1994 (Meredith
and King, 2005). The warming is likely to have resulted from
large regional atmospheric warming (Vaughan et al., 2003) and
reduced winter sea ice observed in this region.
In the Ross Sea and near the Ross Ice Shelf, signicant
decreases in salinity of 0.003 psu yr–1 (and density decreases)
over the last four decades (Jacobs et al., 2002) have been
observed. Downstream of the Ross Ice Shelf in the Australian-
Antarctic Basin, AABW has also cooled and freshened (Aoki
et al., 2005b). These observed decreases are signicantly
greater than earlier reports of AABW variability (Whitworth,
2002) and suggest that changes in the antarctic shelf waters can
be quite quickly communicated to deep waters. Jacobs et al.
(2002) concluded that the freshening appears to have resulted
from a combination of factors including increased precipitation,
reduced sea ice production and increased melting of the West
Antarctic Ice Sheet.
5.3.6 Relation of Regional to Global Changes Changes in Global Water Mass Properties
The regional analyses described in the previous sections
have global organisation, as described partially in Section 5.3.1
(Figure 5.6), and as reected in the global trend analyses in
Section 5.2. The data sets used for the largest-scale descriptions
over the last 30 to 50 years are reliable; different types of data
and widely varying methods yield similar results, increasing
condence in the reality of the changes found in both the global
and regional analyses.
The regional and global analyses of ocean warming generally
show a pattern of increased ocean temperature in the regions of
very thick surface mixed layer (mode water) formation. This is
clearest in the North Atlantic and North Pacic and in all sectors
of the Southern Ocean (Figure 5.3). There are also regions of
decreased ocean temperature in both the global and regional
analyses in parts of the subpolar and equatorial regions.
Both the global and regional analyses show long-term
freshening in the subpolar waters in the North Atlantic and
North Pacic and a salinity increase in the upper ocean (<100 m
deep) at low to mid-latitudes. This is consistent with an increase
in the atmospheric hydrological cycle over the oceans and could
result in changes in ocean advection (Section 5.3.2). In the North
Atlantic, the subpolar freshening occurred throughout the entire
water column, from the 1960s to the mid-1990s (Figure 5.5 and
Figure 5.6a). Increased salinity and temperature in the upper
water column in the subpolar North Atlantic after 1994 are not
apparent from the linear trend applied to the full time series
in Figure 5.5, but are clear in all regional time series (Section
5.3.2). Freshening in the North Pacic subpolar gyre north of
45°N is apparent in both regional analyses (Section and
global analyses (Figure 5.5). Freshening of intermediate depth
waters (>300 m) from Southern Ocean sources (Section 5.3.5)
is apparent in both the global and regional analyses (e.g., Figure
5.5 World).
Many of the observed changes in the temperature and
salinity elds have been linked to atmospheric forcing through
correlations with atmospheric indices associated with the NAO,
PDO and SAM. Indeed, most of the few time series of ocean
measurements or repeat measurements of long sections (see
Sections 5.3.2 and 5.3.4) show evidence of decadal variability.
Because of the long time scales of these natural climate patterns,
it is difcult to discern if observed decadal oceanic variability
is natural or a climate change signal; indeed, changes in
these natural patterns themselves might be related to climate
change. In the North Atlantic, freshening at high latitudes and
increased evaporation at subtropical latitudes prior to the mid-
1990s might have been associated with an increasing NAO
index, and the reversal towards higher salinity at high latitudes
thereafter with a decreasing NAO index after 1990 (see Figure
3.31). Likewise in the Pacic, freshening at high latitudes and
increased evaporation in the subtropics, cooling in the central
North Pacic, warming in the eastern and tropical Pacic and
reduced ventilation in the Kuroshio region, Japan and Okhotsk
Seas could be associated with the extended positive phase of
the PDO. The few detection and attribution studies of ocean
changes are discussed in Section 9.5.1.
At a global scale, the observed long-term patterns of zonal
temperature and salinity changes tend to be approximately
symmetric around the equator (Figure 5.6) and occur
simultaneously in different ocean basins (Figures 5.3 and 5.5).
The scale of these patterns, which extends beyond the regions
of inuence normally associated with the NAO, PDO and SAM,
suggests that these coherent changes between both hemispheres
are associated with a global phenomenon. Consistency with the Large-Scale Ocean
The observed changes are broadly consistent with scientic
understanding of the circulation of the global oceans. The North
Atlantic and antarctic regions, where the oceans ventilate the
deep waters over short time scales (<50 years), show strong
evidence of change over the instrumental record. For example,
the North Atlantic shows evidence of a deep warming and
freshening. There is evidence of change in the Southern Ocean
bottom waters consistent with the sinking of fresher antarctic
shelf waters. Deep waters that are far from the North Atlantic
and Antarctic, remote from interaction with the atmosphere,
and with replenishment rates that are long compared with the
instrumental record, typically show no signicant changes.
Mode waters, key global water masses found in every ocean
basin equatorward of major oceanic frontal systems or
separated boundary currents, have a relatively rapid formation
and ventilation rate (<20 years) and provide a pathway for heat
(and salinity) to be transported into the main subtropical gyres
of the global oceans as observed.
Chapter 5 Observations: Oceanic Climate Change and Sea Level
5.4 Ocean Biogeochemical Changes
5.4.1 Introduction
The observed increase in atmospheric carbon dioxide
(CO2; see Chapter 2) and the observed changes in the physical
properties of the ocean reported in this chapter can affect
marine biogeochemical cycles (here mainly carbon, oxygen,
and nutrients). The increase in atmospheric CO2 causes
additional CO2 to dissolve in the ocean. Changes in temperature
and salinity affect the solubility and chemical equilibration
of gases. Changes in circulation affect the supply of carbon
and nutrients from below, the ventilation of oxygen-depleted
waters and the downward penetration of anthropogenic
carbon. The combined physical and biogeochemical changes
also affect biological activity, with further consequences for
the biogeochemical cycles.
The increase in surface ocean CO2 has consequences for the
chemical equilibrium of the ocean. As CO2 increases, surface
waters become more acidic and the concentration of carbonate
ions decreases. This change in chemical equilibrium causes
a reduction of the capacity of the ocean to take up additional
CO2. However, the response of marine organisms to ocean
acidication is poorly known and could cause further changes
in the marine carbon cycle with consequences that are difcult
to estimate (see Section 7.3.4 and Chapter 4 of the Working
Group II contribution to the IPCC Fourth Assessment Report).
Dissolved oxygen (O2) in the ocean is affected by the same
physical processes that affect CO2, but in contrast to CO2, O2
is not affected by changes in its atmospheric concentration
(which are only of the order of 10–4 of its mean concentration).
Changes in oceanic O2 concentration thus provide information
on the changes in the physical or biological processes that occur
within the ocean, such as ventilation (here used to describe the
rate of renewal of thermocline waters), mode water formation,
upwelling or biological export and respiration. Furthermore,
changes in the oceanic O2 content are needed to estimate the
CO2 budget from atmospheric O2/molecular nitrogen (N2)
ratio measurements. However, the method currently estimates
the change in air-sea uxes of O2 indirectly based on heat ux
changes (see Section 7.3.2).
This section reports observed changes in biogeochemical
cycles and assesses their consistency with observed changes in
physical properties. Changes in oceanic nitrous oxide (N2O) and
methane (CH4) have not been assessed because of the lack of
large-scale observations. Observations of the mean uxes of N2O
and CH4 (including CH4 hydrates) are discussed in Chapter 7.
5.4.2 Carbon Total Change in Dissolved Inorganic Carbon and
Air-Sea Carbon Dioxide Flux
Direct observations of oceanic dissolved inorganic carbon
(DIC; i.e., the sum of CO2 plus carbonate and bicarbonate)
reect changes in both the natural carbon cycle and the uptake
of anthropogenic CO2 from the atmosphere. Links between the
main modes of climate variability and the marine carbon cycle
have been observed on interannual time scales in several regions
of the world (see Section for quantitative estimates). In
the equatorial Pacic, the reduced upwelling associated with El
Niño events decreases the regional outgas of natural CO2 to the
atmosphere (Feely et al., 1999). In the subtropical North Atlantic,
reduced mode water formation and reduced deep winter mixing
during the positive NAO phase increase the storage of carbon in
the intermediate ocean (Bates et al., 2002). These observations
show that variability in the content of natural DIC in the ocean
has occurred in association with climate variability.
Longer observations exist for the partial pressure of CO2
(pCO2) at the surface only. Over more than two decades, the
oceanic pCO2 increase has generally followed the atmospheric
CO2 within the given uncertainty, although regional differences
have been observed (Feely et al., 1999; Takahashi et al., 2006).
The three stations with the longest time series, all in the northern
subtropics, show pCO2 increases at a rate varying between
1.6 and 1.9 μatm yr–1 (Figure 5.9), indistinguishable from the
atmospheric increase of 1.5 to 1.9 μatm yr–1. Variability on the
order of 20 μatm over periods of ve years was observed in
the three time series, as well as in other data sets, and has been
associated with regional changes in the natural carbon cycle
driven by changes in ocean circulation and by climate variability
(Gruber et al., 2002; Dore et al., 2003) or with variations in
biological activity (Lefèvre et al., 2004).
Direct surface pCO2 observations have been used to compute
a global air-sea CO2 ux of 1.6 ± 1 GtC yr–1 for the year 1995
(Takahashi et al., 2002; Section, Figure 7.8). It is not
yet possible to detect large-scale changes in the global air-sea
CO2 ux from direct observations because of the large inuence
of climate variability. However, estimates from inverse methods
of the air-sea CO2 ux from the spatio-temporal distribution of
atmospheric CO2 suggest that the global air-sea CO2 ux increased
by 0.1 to 0.6 GtC yr–1 between the 1980s and 1990s, consistent
with results from ocean models (Le Quéré et al., 2003). Anthropogenic Carbon Change
The recent uptake of anthropogenic carbon in the ocean is
well constrained by observations to a decadal mean of 2.2 ±
0.4 GtC yr–1 for the 1990s (see Section 7.3.2, Table 7.1). The
uptake of anthropogenic carbon over longer time scales can be
estimated from oceanic measurements. Changes in DIC between
two time periods reect the anthropogenic carbon uptake plus
the changes in DIC concentration due to changes in water
masses and biological activity. To estimate the contribution
of anthropogenic carbon alone, several corrections must be
applied. From observed DIC changes between surveys in the
1970s and the 1990s, an increase in anthropogenic carbon has
been inferred down to depths of 1,100 m in the North Pacic
(Peng et al., 2003; Sabine et al., 2004a), 200 to 1,200 m in
the Indian Ocean (Peng et al., 1998; Sabine et al., 1999) and
1,900 m in the Southern Ocean (McNeil et al., 2003).
Observations: Oceanic Climate Change and Sea Level Chapter 5
An indirect method was used to estimate
anthropogenic carbon from observations made at a
single time period based on well-known processes that
control the distribution of natural DIC in the ocean.
The method corrects the observed DIC concentration
for organic matter decomposition and dissolution of
carbonate minerals, and removes an estimate of the
DIC concentration of the water when it was last in
contact with the atmosphere (Gruber et al., 1996).
With this method, a global DIC increase of 118 ±
19 GtC between pre-industrial times (roughly 1750)
and 1994 has been estimated, using 9,618 proles
from the 1990s (Sabine et al., 2004b; see Figure 5.10).
The uncertainty of ±19 GtC in this estimate is based
on uncertainties in the anthropogenic DIC estimates
and mapping errors, which have characteristics of
random error, and on an estimate of potential biases,
which are not necessarily centred on the mean value.
Potential biases of up to 7% in the technique have been
identied, mostly caused by assumptions about the
time evolution of CO2, the age or the identication of
water masses (Matsumoto and Gruber, 2005), and the
recent changes in surface warming and stratication
(Keeling, 2005). Potential biases from assumptions
of constant carbon and nutrient uptake ratios for
biological activity have not been assessed. While the
magnitude and direction of all potential biases are
not yet clear, the given uncertainty of ±16% appears
realistic compared to the biases already identied.
Because of the limited rate of vertical transport in the ocean,
more than half of the anthropogenic carbon can still be found
in the upper 400 m, and it is undetectable in most of the deep
ocean (Figure 5.11). The vertical penetration of anthropogenic
carbon is consistent with the DIC changes observed between
two cruises (Peng et al., 1998, 2003). Anthropogenic carbon
has penetrated deeper in the North Atlantic and subantarctic
Southern Ocean compared to other basins, due to a combination
of: i) high surface alkalinity (in the Atlantic) which favours the
uptake of CO2, and ii) more active vertical exchanges caused
by intense winter mixing and by the formation of deep waters
(Sabine et al., 2004b). The deeper penetration of anthropogenic
carbon in these regions is consistent with similar features
observed in the oceanic distribution of chlorouorocarbons
(CFCs) of atmospheric origin (Willey et al., 2004), conrming
that it takes decades to many centuries to transport carbon from
the surface into the thermocline and the deep ocean. Deeper
penetration in the North Atlantic and subantarctic Southern
Ocean is also observed in the changes in heat content shown in
Figure 5.3. The large storage of anthropogenic carbon observed
in the subtropical gyres is caused by the lateral transport of
carbon from the region of mode water formation towards the
lower latitudes (Figure 5.10).
The fraction of the net CO2 emissions taken up by the ocean
(the uptake fraction) was possibly lower during 1980 to 2005
(37% ± 7%) compared to 1750 to 1994 (42% ± 7%); however
the uncertainty in the estimates is larger than the difference
between the estimates (Table 5.1). The net CO2 emissions
Table 5.1. Fraction of CO2 emissions taken up by the ocean for different time periods.
Time Period Oceanic Increase (GtC) Net CO2 Emissionsa (GtC) Uptake Fraction (%) Reference
1750–1994 118 ± 19 283 ± 19 42 ± 7 Sabine et al., 2004b
1980–2005b 53 ± 9 143 ± 10 37 ± 7 Chapter 7c
Figure 5.9. Changes in surface oceanic pCO2 (left; in μatm) and pH (right) from three time
series stations: Blue: European Station for Time-series in the Ocean (ESTOC, 29°N, 15°W;
Gonzalez-Dávila et al., 2003); green: Hawaii Ocean Time-Series (HOT, 23°N, 158°W; Dore et
al., 2003); red: Bermuda Atlantic Time-series Study (BATS, 31/32°N, 64°W; Bates et al., 2002;
Gruber et al., 2002). Values of pCO2 and pH were calculated from DIC and alkalinity at HOT and
BATS; pH was directly measured at ESTOC and pCO2 was calculated from pH and alkalinity. The
mean seasonal cycle was removed from all data. The thick black line is smoothed and does not
contain variability less than 0.5 years period.
a Sum of emissions from fossil fuel burning, cement production, land use change and the terrestrial biosphere response.
b The longest possible time period was used for the recent decades to minimise the effect of the variability in atmospheric CO2.
c Sum of the estimates for the 1980s, 1990s and 2000 to 2005 from Table 7.1.
Chapter 5 Observations: Oceanic Climate Change and Sea Level
include all emissions that have an inuence on the atmospheric
CO2 concentration (i.e., emissions from fossil fuel burning,
cement production, land use change and the terrestrial biosphere
response). It is equivalent to the sum of the atmospheric and
oceanic CO2 increase. Because the atmospheric CO2 is well
constrained by observations, the uncertainty in the net CO2
emissions is nearly equal to the uncertainty in the oceanic
CO2 increase. The decrease in oceanic uptake fraction would
be consistent with the understanding that the ocean CO2 sink
is limited by the transport rate of anthropogenic carbon from
the surface to the deep ocean, and also with the nonlinearity in
carbon chemistry that reduces the CO2 uptake capacity of water
as its CO2 concentration increases (Sarmiento et al., 1995). OceanAcidicationbyCarbonDioxide
The uptake of anthropogenic carbon by the ocean changes
the chemical equilibrium of the ocean. Dissolved CO2 forms
a weak acid.1 As CO2 increases, pH decreases, that is, the
ocean becomes more acidic. Ocean pH can be computed from
measurements of DIC and alkalinity. A decrease in surface pH
of 0.1 over the global ocean was calculated from the estimated
uptake of anthropogenic carbon between 1750 and 1994 (Sabine
et al., 2004b; Raven et al., 2005), with the lowest decrease (0.06)
in the tropics and subtropics, and the highest decrease (0.12)
at high latitudes, consistent with the lower buffer capacity of
the high latitudes compared to the low latitudes. The mean
pH of surface waters ranges between 7.9 and 8.3 in the open
Figure 5.10. Column inventory of anthropogenic carbon (mol m–2) as of 1994 from Sabine et al. (2004b). Anthropogenic carbon is estimated indirectly by correcting the
measured DIC for the contributions of organic matter decomposition and dissolution of carbonate minerals, and taking into account the DIC concentration the water had in the
pre-industrial ocean when it was last in contact with the atmosphere. The global inventory of anthropogenic carbon taken up by the ocean between 1750 and 1994 is estimated
to be 118 ± 19 GtC.
Figure 5.11. Mean concentration of anthropogenic carbon as of 1994 in μmol
kg–1 from Sabine et al. (2004b) averaged over (a) the Pacific and Indian Oceans and
(b) the Atlantic Ocean. The calculation of anthropogenic carbon is described in the
caption of Figure 5.10 and in the text (Section 5.4).
1 Acidity is a measure of the concentration of H+ ions and is reported in pH units, where pH = –log10(H+). A pH decrease of 1 unit means a 10-fold increase in the concentration of
H+, or acidity.
Observations: Oceanic Climate Change and Sea Level Chapter 5
ocean, so the ocean remains alkaline (pH > 7) even after these
decreases. For comparison, pH was higher by 0.1 unit during
glaciations, and there is no evidence of pH values more than 0.6
units below the pre-industrial pH during the past 300 million
years (Caldeira and Wickett, 2003). A decrease in ocean pH of
0.1 units corresponds to a 30% increase in the concentration
of H+ in seawater, assuming that alkalinity and temperature
remain constant. Changes in surface temperature may have
induced an additional decrease in pH of <0.01. The calculated
anthropogenic impact on pH is consistent with results from
time series stations where a decrease in pH of 0.02 per decade
was observed (Figure 5.9). Results from time series stations
include not only the increase in anthropogenic carbon, but also
other changes due to local physical and biological variability.
The consequences of changes in pH on marine organisms are
poorly known (see Section 7.3.4 and Box 7.3). Change in Carbonate Species
The uptake of anthropogenic carbon occurs through the
injection of CO2 and causes a shift in the distribution of
carbon species (i.e., the balance between CO2, carbonate and
bicarbonate). The availability of carbonate is particularly
important because it controls the maximum amount of CO2 that
the ocean is able to absorb. Marine organisms use carbonate
to produce shells of calcite and aragonite (both consisting of
calcium carbonate; CaCO3). Currently, the surface ocean is
super-saturated with respect to both calcite and aragonite, but
undersaturated below a depth called the ‘saturation horizon’.
The undersaturation starts at a depth varying between 200 m in
parts of the high-latitude and the Indian Ocean and 3,500 m in
the Atlantic. Calcium carbonate dissolves either when it sinks
below the calcite or aragonite saturation horizons or under the
action of biological activity.
Shoaling of the aragonite saturation horizon has been
observed in all ocean basins based on alkalinity, DIC and oxygen
measurements (Feely and Chen, 1982; Feely et al., 2002; Sabine
et al., 2002; Sarma et al., 2002). The amplitude and direction
of the signal was everywhere consistent with the uptake of
anthropogenic carbon, with potentially smaller contributions
from changes in circulation, temperature and biology. Feely et
al. (2004) calculated that the uptake of anthropogenic carbon
alone has caused a shoaling of the aragonite saturation horizon
between 1750 and 1994 by 30 to 200 m in the eastern Atlantic
(50°S–15°N), the North Pacic and the North Indian Ocean,
and a shoaling of the calcite saturation horizon by 40 to
100 m in the Pacic (north of 20°N). This calculation is based
on the anthropogenic DIC increase estimated by Sabine et al.
(2004a), on a global compilation of biogeochemical data and on
carbonate chemistry equations. Furthermore, an increase in total
alkalinity (primarily controlled by carbonate and bicarbonate)
at the depth of the aragonite saturation horizon between 1970
and 1990 has been reported (Sarma et al., 2002). These results
are consistent with the calculated increase in CaCO3 dissolution
as a result of the shoaling of the aragonite saturation horizon,
but with large uncertainty. Carbonate decreases at high latitudes
and particularly in the Southern Ocean may have consequences
for marine ecosystems because the current saturation horizon is
closer to the surface than in other basins (Orr et al., 2005; see
Section 7.3.4).
5.4.3 Oxygen
In the thermocline (~100 to 1,000 m), a decrease in the O2
concentration has been observed between about the early 1970s
and the late 1990s or later in several repeated hydrographic
sections in the North and South Pacic, North Atlantic, and
Southern Indian Oceans (Figure 5.12; see summary table in
Emerson et al., 2004, and Section 5.3). Section 5.3 reports
on a number of O2 decreases that t the overall message
of Section 5.4. The reported O2 decreases range from 0.1 to
6 μmol kg–1 yr–1, superposed on decadal variations of ±2
μmol kg–1 yr–1 (Ono et al., 2001; Andreev and Watanabe, 2002).
In all published studies, the observed O2 decrease appeared to
be driven primarily by changes in ocean circulation, and less by
changes in the rate of O2 demand from downward settling of
organic matter. A few studies have quantied the contribution
of the change in ocean circulation using estimates of changes in
apparent CFC ages (Doney et al., 1998; Watanabe et al., 2001;
Mecking et al., 2006). In nearly all cases, the decrease in O2
could entirely be accounted for by the increased apparent CFC
age that resulted from reduced rate of renewal of intermediate
waters. Changes in biological processes were only signicant at
the coast of California and may result from assumptions in the
method (Mecking et al., 2006).
It is unclear whether the recent changes in O2 are indicative
of trends or of variability. Recent data in the Indian Ocean
have shown a reversal of the O2 decrease between 1987 and
2002 in the South Indian Ocean of similar amplitude to the
decrease observed during the previous decades (McDonagh et
al., 2005). Variability has been observed on decadal time scales
in the North Atlantic large enough to mask any potential trends
(Johnson and Gruber, 2007).
In the upper 100 m of the global ocean surface, decadal
variations of ±0.5 μmol kg–1 in O2 concentration were observed
for the period 1956 to 1998 based on a global analysis of
530,000 oxygen proles, with no clear trends (Garcia et al.,
2005). However, the near-surface changes in O2 concentration
are difcult to interpret. They can be caused by changes in
biological activity, by changes in the physical transport of
O2 from intermediate waters or by changes in temperature
and salinity. Because there is less condence in the early
measurements and the reported changes cannot be explained
by known processes, it cannot be said whether the absence of a
long-term trend in surface O2 is realistic or not.
5.4.4 Nutrients
Changes in nutrient concentrations can provide information
on changes in the physical and biological processes that affect
the carbon cycle and could potentially be used as indicators
for large-scale changes in marine biology. However, only a
Chapter 5 Observations: Oceanic Climate Change and Sea Level
few studies reported decadal changes in inorganic nutrient
concentrations. In the North Pacic, the concentration of nitrate
plus nitrite (N) and phosphate decreased at the surface (Freeland
et al., 1997; Watanabe et al., 2005) and increased below the
surface (Emerson et al., 2001; Ono et al., 2001; Keller et al.,
2002) in the past two decades. Nutrient changes were observed
in the deep ocean of all basins but no clear pattern emerges
from available observations. Pahlow and Riebesell (2000)
found changes in the ratio of nutrients in the North Pacic and
Atlantic Oceans, and no signicant changes in the South Pacic.
In the North Pacic, Keller et al. (2002) observed a decrease in
N associated with the increase in O2 between 1970 and 1990
at 1,050 m, opposite to the results of Pahlow and Riebesell’s
longer study. Using the same data set extended to the world,
large regional changes in nutrient ratios were observed (Li and
Peng, 2002) but no consistent basin-scale patterns. Uncertainties
in deep ocean nutrient observations may be responsible for
the lack of coherence in the nutrient changes. Sources of
inaccuracy include the limited number of observations and the
lack of compatibility between measurements from different
laboratories at different times.
In some cases, the observed trends in nutrients can be
explained by either a change in thermocline ventilation or
a change in biological activity (Pahlow and Riebesell, 2000;
Emerson et al., 2001), but in other cases are mostly consistent
with a reduction in thermocline ventilation (Freeland et al.,
1997; Ono et al., 2001; Watanabe et al., 2005). Thus, all of the
reported trends are consistent with a physical explanation of
the observed changes, although changes in biological activity
cannot be ruled out.
The concentration of surface nutrients can also be inuenced
by surface mixing, as a reduction in mixing leads to a decreased
concentration of surface nutrients. The observed changes in
surface temperature and salinity (see Sections 5.2.3 and 5.3) are
indicative of changes in the surface mixing (see Section
In most of the Pacic Ocean, surface warming and freshening
act in the same direction and contribute to reduced mixing
(Figures 5.2 and 5.5), consistent with regional observations
(Freeland et al., 1997; Watanabe et al., 2005). In the Atlantic
and Indian Oceans, temperature and salinity trends generally
act in opposite directions and changes in mixing have not been
quantied regionally.
Figure 5.12. Changes in oxygen concentration (μmol kg–1) along two sections in the North Pacific (see map, bottom panel). Top left panel: Difference (1999 minus 1985)
along 47°N. Top right panel: Difference (1997 minus 1984) at 152°W. Blue colours indicate a decrease and yellow colours indicate an increase in oxygen over time. The differ-
ences were calculated using density as the vertical coordinate. After Deutsch et al. (2005).
Observations: Oceanic Climate Change and Sea Level Chapter 5
in ocean physics inuence natural biogeochemical cycles, and
thus the cycles of O2 and CO2 are likely to undergo changes if
ocean circulation changes persist in the future.
5.5 Changes in Sea Level
5.5.1 Introductory Remarks
Present-day sea level change is of considerable interest
because of its potential impact on human populations living in
coastal regions and on islands. This section focuses on global
and regional sea level variations, over time spans ranging from
the last decade to the past century; a brief discussion of sea level
change in previous centuries is given in Section Changes
over previous millennia are discussed in Section 6.4.3.
Processes in several nonlinearly coupled components of the
Earth system contribute to sea level change, and understanding
these processes is therefore a highly interdisciplinary endeavour.
On decadal and longer time scales, global mean sea level change
results from two major processes, mostly related to recent
climate change, that alter the volume of water in the global
ocean: i) thermal expansion (Section 5.5.3), and ii) the exchange
of water between oceans and other reservoirs (glaciers and ice
caps, ice sheets, other land water reservoirs - including through
anthropogenic change in land hydrology, and the atmosphere;
Section 5.5.5). All these processes cause geographically non-
uniform sea level change (Section 5.5.4) as well as changes in
the global mean; some oceanographic factors (e.g., changes in
ocean circulation or atmospheric pressure) also affect sea level
at the regional scale, while contributing negligibly to changes
in the global mean. Vertical land movements such as resulting
from glacial isostatic adjustment (GIA), tectonics, subsidence
and sedimentation inuence local sea level measurements but
do not alter ocean water volume; nonetheless, they affect global
mean sea level through their alteration of the shape and hence
the volume of the ocean basins containing the water.
Measurements of present-day sea level change rely on two
different techniques: tide gauges and satellite altimetry (Section
5.5.2). Tide gauges provide sea level variations with respect to
the land on which they lie. To extract the signal of sea level
change due to ocean water volume and other oceanographic
change, land motions need to be removed from the tide gauge
measurement. Land motions related to GIA can be simulated
in global geodynamic models. The estimation of other land
motions is not generally possible unless there are adequate
nearby geodetic or geological data, which is usually not the case.
However, careful selection of tide gauge sites such that records
reecting major tectonic activity are rejected, and averaging
over all selected gauges, results in a small uncertainty for global
sea level estimates (Appendix 5.A.4). Sea level change based
on satellite altimetry is measured with respect to the Earth’s
centre of mass, and thus is not distorted by land motions, except
for a small component due to large-scale deformation of ocean
basins from GIA.
5.4.5 Biological Changes Relevant to Ocean
Changes in biological activity are an important part of the
carbon cycle but are difcult to quantify at the global scale.
Marine export production (the fraction of primary production
that is not respired at the ocean surface and thus sinks to depth)
is the biological process that has the largest inuence on
element cycles. There are no global observations on changes
in export production or respiration. However, estimates of
changes in primary production provide partial information.
A reduction in global oceanic primary production by about
6% between the early 1980s and the late 1990s was estimated
based on the comparison of chlorophyll data from two satellites
(Gregg et al., 2003). The errors in this estimate are potentially
large because it is based on the comparison of data from two
different sensors. Nevertheless, a change in biological uxes of
this order of magnitude is plausible considering that biological
production is controlled primarily by nutrient input from
intermediate waters, and that a decrease in intermediate water
renewal has been observed during that period as indicated by
the decrease in O2. Shifts and trends in plankton biomass have
been observed for instance in the North Atlantic (Beaugrand
and Reid, 2003), the North Pacic (Karl, 1999; Chavez et
al., 2003) and in the Southern Indian Ocean (Hirawake et al.,
2005), but the spatial and temporal coverage is limited. The
potential impacts of changes in marine ecosystems or dissolved
organic matter on climate are discussed in Section 7.3.4, and
the impact of climate on marine ecosystems in Chapter 4 of
the Working Group II contribution to the IPCC Fourth
Assessment Report.
5.4.6 Consistency with Physical Changes
It is clearly established that climate variability affects the
oceanic content of natural and anthropogenic DIC and the air-
sea ux of CO2, although the amplitude and physical processes
responsible for the changes are less well known. Variability
in the marine carbon cycle has been observed in response
to physical changes associated with the dominant modes of
climate variability such as El Niño events and the PDO (Feely
et al., 1999; Takahashi et al., 2006), and the NAO (Bates et
al., 2002; Johnson and Gruber, 2007). The regional patterns of
anthropogenic CO2 storage are consistent with those of CFCs
and with changes in heat content. The observed trends in CO2,
DIC, pH and carbonate species can be primarily explained by
the response of the ocean to the increase in atmospheric CO2.
Large-scale changes in the O2 content of the thermocline
have been observed between the 1970s and the late 1990s.
These changes are everywhere consistent with the local changes
in ocean ventilation as identied either by changes in density
gradients or by changes in apparent CFC ages. Nevertheless,
an inuence of changes in marine biology cannot be ruled out.
The available data are insufcient to say if the changes in O2
are caused by natural variability or are trends that are likely to
persist in the future, but they do indicate that large-scale changes
Chapter 5 Observations: Oceanic Climate Change and Sea Level
Frequently Asked Question 5.1
Is Sea Level Rising?
Yes, there is strong evidence that global sea level gradually
rose in the 20th century and is currently rising at an increased
rate, after a period of little change between AD 0 and AD 1900.
Sea level is projected to rise at an even greater rate in this century.
The two major causes of global sea level rise are thermal expan-
sion of the oceans (water expands as it warms) and the loss of
land-based ice due to increased melting.
Global sea level rose by about 120 m during the several mil-
lennia that followed the end of the last ice age (approximately
21,000 years ago), and stabilised between 3,000 and 2,000 years
ago. Sea level indicators suggest that global sea level did not
change significantly from then until the late 19th century. The
instrumental record of modern sea level change shows evidence
for onset of sea level rise during the 19th century. Estimates for
the 20th century show that global average sea level rose at a rate
of about 1.7 mm yr–1.
Satellite observations available since the early 1990s provide
more accurate sea level data with nearly global coverage. This
decade-long satellite altimetry data set shows that since 1993, sea
level has been rising at a rate of around 3 mm yr–1, significantly
higher than the average during the previous half century. Coastal
tide gauge measurements confirm this observation, and indicate
that similar rates have occurred in some earlier decades.
In agreement with climate models, satellite data and hydro-
graphic observations show that sea level is not rising uniformly
around the world. In some regions, rates are up to several times the
global mean rise, while in other regions sea level is falling. Sub-
stantial spatial variation in rates of sea level change is also inferred
from hydrographic observations. Spatial variability of the rates of
sea level rise is mostly due to non-uniform changes in temperature
and salinity and related to changes in the ocean circulation.
Near-global ocean temperature data sets made available in
recent years allow a direct calculation of thermal expansion. It
is believed that on average, over the period from 1961 to 2003,
thermal expansion contributed about one-quarter of the observed
sea level rise, while melting of land ice accounted for less than
half. Thus, the full magnitude of the observed sea level rise during
that period was not satisfactorily explained by those data sets, as
reported in the IPCC Third Assessment Report.
During recent years (1993–2003), for which the observing
system is much better, thermal expansion and melting of land
ice each account for about half of the observed sea level rise,
although there is some uncertainty in the estimates.
The reasonable agreement in recent years between the observed
rate of sea level rise and the sum of thermal expansion and loss of
land ice suggests an upper limit for the magnitude of change in
land-based water storage, which is relatively poorly known. Mod-
el results suggest no net trend in the storage of water over land
due to climate-driven changes but there are large interannual and
decadal fluctuations. However, for the recent period 1993 to 2003,
the small discrepancy between observed sea level rise and the sum
of known contributions might be due to unquantified human-
induced processes (e.g., groundwater extraction, impoundment in
reservoirs, wetland drainage and deforestation).
Global sea level is projected to rise during the 21st century at
a greater rate than during 1961 to 2003. Under the IPCC Special
Report on Emission Scenarios (SRES) A1B scenario by the mid-
2090s, for instance, global sea level reaches 0.22 to 0.44 m above
1990 levels, and is rising at about 4 mm yr–1. As in the past, sea
level change in the future will not be geographically uniform,
with regional sea level change varying within about ±0.15 m of
the mean in a typical model projection. Thermal expansion is pro-
jected to contribute more than half of the average rise, but land
ice will lose mass increasingly rapidly as the century progresses.
An important uncertainty relates to whether discharge of ice from
the ice sheets will continue to increase as a consequence of accel-
erated ice flow, as has been observed in recent years. This would
add to the amount of sea level rise, but quantitative projections of
how much it would add cannot be made with confidence, owing
to limited understanding of the relevant processes.
Figure 1 shows the evolution of global mean sea level in
the past and as projected for the 21st century for the SRES A1B
FAQ 5.1, Figure 1. Time series of global mean sea level (deviation from the
1980-1999 mean) in the past and as projected for the future. For the period before
1870, global measurements of sea level are not available. The grey shading shows
the uncertainty in the estimated long-term rate of sea level change (Section 6.4.3).
The red line is a reconstruction of global mean sea level from tide gauges (Section, and the red shading denotes the range of variations from a smooth curve.
The green line shows global mean sea level observed from satellite altimetry. The
blue shading represents the range of model projections for the SRES A1B scenario
for the 21st century, relative to the 1980 to 1999 mean, and has been calculated
independently from the observations. Beyond 2100, the projections are increasingly
dependent on the emissions scenario (see Chapter 10 for a discussion of sea level
rise projections for other scenarios considered in this report). Over many centuries or
millennia, sea level could rise by several metres (Section 10.7.4).
Observations: Oceanic Climate Change and Sea Level Chapter 5
The TAR chapter on sea level change provided
estimates of climate and other anthropogenic
contributions to 20th-century sea level rise, based
mostly on models (Church et al., 2001). The
sum of these contributions ranged from –0.8 to
2.2 mm yr–1, with a mean value of 0.7 mm yr–1,
and a large part of this uncertainty was due to the
lack of information on anthropogenic land water
change. For observed 20th-century sea level rise,
based on tide gauge records, Church et al. (2001)
adopted as a best estimate a value in the range of 1
to 2 mm yr–1, which was more than twice as large as
the TAR’s estimate of climate-related contributions.
It thus appeared that either the processes causing
sea level rise had been underestimated or the rate
of sea level rise observed with tide gauges was
biased towards higher values.
Since the TAR, a number of new results have
been published. The global coverage of satellite
altimetry since the early 1990s (TOPography
EXperiment (TOPEX)/Poseidon and Jason) has
improved the estimate of global sea level rise and
has revealed the complex geographical patterns
of sea level change in open oceans. Near-global
ocean temperature data for the last 50 years have
been recently made available, allowing the rst observationally
based estimate of the thermal expansion contribution to sea
level rise in past decades. For recent years, better estimates of
the land ice contribution to sea level are available from various
observations of glaciers, ice caps and ice sheets.
In this section, we summarise the current knowledge of
present-day sea level rise. The observational results are assessed,
followed by our current interpretation of these observations in
terms of climate change and other processes, and ending with a
discussion of the sea level budget (Section 5.5.6).
5.5.2 Observations of Sea Level Changes 20th-Century Sea Level Rise from Tide Gauges
Table 11.9 of the TAR listed several estimates for global and
regional 20th-century sea level trends based on the Permanent
Service for Mean Sea Level (PSMSL) data set (Woodworth
and Player, 2003). The concerns about geographical bias in
the PSMSL data set remain, with most long sea level records
stemming from the NH, and most from continental coastlines
rather than ocean interiors. Based on a small number (~25) of
high-quality tide gauge records from stable land regions, the
rate of sea level rise has been estimated as 1.8 mm yr–1 for the
past 70 years (Douglas, 2001; Peltier, 2001), and Miller and
Douglas (2004) nd a range of 1.5 to 2.0 mm yr–1 for the 20th
century from 9 stable tide gauge sites. Holgate and Woodworth
(2004) estimated a rate of 1.7 ± 0.4 mm yr–1 sea level change
averaged along the global coastline during the period 1948 to
2002, based on data from 177 stations divided into 13 regions.
Church et al. (2004) (discussed further below) determined
a global rise of 1.8 ± 0.3 mm yr–1 during 1950 to 2000, and
Church and White (2006) determined a change of 1.7 ±
0.3 mm yr–1 for the 20th century. Changes in global sea level
as derived from analyses of tide gauges are displayed in Figure
5.13. Considering the above results, and allowing for the
ongoing higher trend in recent years shown by altimetry (see
Section, we assess the rate for 1961 to 2003 as 1.8 ±
0.5 mm yr–1 and for the 20th century as 1.7 ± 0.5 mm yr–1.
While the recently published estimates of sea level rise over
the last decades remain within the range of the TAR values
(i.e., 1–2 mm yr–1), there is an increasing opinion that the best
estimate lies closer to 2 mm yr–1 than to 1 mm yr–1. The lower
bound reported in the TAR resulted from local and regional
studies; local and regional rates may differ from the global
mean, as discussed below (see Section
A critical issue concerns how the records are adjusted for
vertical movements of the land upon which the tide gauges
are located and of the oceans. Trends in tide gauge records are
corrected for GIA using models, but not for other land motions.
The GIA correction ranges from about 1 mm yr–1 (or more) near
to former ice sheets to a few tenths of a millimetre per year in
the far eld (e.g., Peltier, 2001); the error in tide-gauge based
global average sea level change resulting from GIA is assessed
as 0.15 mm yr–1. The TAR mentioned the developing geodetic
technologies (especially the Global Positioning System; GPS)
that hold the promise of measuring rates of vertical land
movement at tide gauges, no matter if those movements are
due to GIA or to other geological processes. Although there
has been some model validation, especially for GIA models,
systematic problems with such techniques, including short data
spans, have yet to be fully resolved.
Figure 5.13. Annual averages of the global mean sea level (mm). The red curve shows reconstructed
sea level fields since 1870 (updated from Church and White, 2006); the blue curve shows coastal tide
gauge measurements since 1950 (from Holgate and Woodworth, 2004) and the black curve is based
on satellite altimetry (Leuliette et al., 2004). The red and blue curves are deviations from their averages
for 1961 to 1990, and the black curve is the deviation from the average of the red curve for the period
1993 to 2001. Error bars show 90% confidence intervals.
Chapter 5 Observations: Oceanic Climate Change and Sea Level Sea Level Change during the Last Decade from
Satellite Altimetry
Since 1992, global mean sea level can be computed at 10-
day intervals by averaging the altimetric measurements from
the TOPEX/Poseidon (T/P) and Jason satellites over the area
of coverage (66°S to 66°N) (Nerem and Mitchum, 2001). Each
10-day estimate of global mean sea level has an accuracy of
approximately 5 mm. Numerous papers on the altimetry results
(see Cazenave and Nerem, 2004, for a review) show a current
rate of sea level rise of 3.1 ± 0.7 mm yr–1 over 1993 to 2003
(Cazenave and Nerem, 2004; Leuliette et al., 2004; Figure
5.14). A signicant fraction of the 3 mm yr–1 rate of change
has been shown to arise from changes in the Southern Ocean
(Cabanes et al., 2001).
The accuracy needed to compute mean sea level change
pushes the altimeter measurement system to its performance
limits, and thus care must be taken to ensure that the instrument
is precisely calibrated (see Appendix 5.A.4.1). The tide gauge
calibration method (Mitchum, 2000) provides diagnoses of
problems in the altimeter instrument, the orbits, the measurement
corrections and ultimately the nal sea level data. Errors in
determining the altimeter instrument drift using the tide gauge
calibration, currently estimated to be about 0.4 mm yr–1, are
almost entirely driven by errors in knowledge of vertical land
motion at the gauges (Mitchum, 2000).
Altimetry-based sea level measurements include variations
in the global ocean basin volume due to GIA. Averaged over the
oceanic regions sampled by the altimeter satellites, this effect
yields a value close to –0.3 mm yr–1 in sea level (Peltier, 2001),
with possible uncertainty of 0.15 mm yr–1. This number is
subtracted from altimetry-derived global mean sea level in order
to obtain the contribution due to ocean (water) volume change.
Altimetry from T/P allows the mapping of the geographical
distribution of sea level change (Figure 5.15a). Although
regional variability in coastal sea level change had been reported
from tide gauge analyses (e.g., Douglas, 1992; Lambeck, 2002),
the global coverage of satellite altimetry provides unambiguous
evidence of non-uniform sea level change in open oceans,
with some regions exhibiting rates of sea level change about
ve times the global mean. For the past decade, sea level rise
shows the highest magnitude in the western Pacic and eastern
Indian oceans, regions that exhibit large interannual variability
associated with ENSO. Except for the Gulf Stream region,
most of the Atlantic Ocean shows sea level rise during the past
decade. Despite the global mean rise, Figure 5.15a shows that
sea level has been dropping in some regions (eastern Pacic
and western Indian Oceans). These spatial patterns likely reect
decadal uctuations rather than long-term trends. Empirical
Orthogonal Functions (EOF) analyses of altimetry-based sea
level maps over 1993 to 2003 show a strong inuence of the
1997–1998 El Niño, with the geographical patterns of the
dominant mode being very similar to those of the sea level trend
map (e.g., Nerem et al., 1999). Reconstructions of Sea Level Change during the
Last 50 Years Based on Satellite Altimetry and
Tide Gauges
Attempts have been made to reconstruct historical sea level
elds by combining the near-global coverage from satellite
altimeter data with the longer but spatially sparse tide gauge
records (Chambers et al., 2002; Church et al., 2004). These sea
level reconstructions use the short altimeter record to determine
the principal EOF of sea level variability, and the tide gauge
data to estimate the evolution of the amplitude of the EOFs over
time. The method assumes that the geographical patterns of
decadal sea level trends can be represented by a superposition
of the patterns of variability that are manifest in interannual
variability. The sea level for the period 1870 to 2000 (Church
and White, 2006) shown in Figure 5.13 is based on this approach.
As a caveat, note that variability on different time scales may
have different characteristic patterns (see Section
The trends in the EOF amplitudes (and the implied global
correlations) allow the reconstruction of a spatially variable
rate of sea level rise. Figure 5.16a (updated from Church et al.,
2004) shows the geographical distribution of linear sea level
trends for 1955 to 2003 based on this reconstruction technique.
Comparison with the altimetry-based trend map for the shorter
period (1993 to 2003) indicates quite different geographical
patterns. These differences mainly arise from thermal expansion
changes through time (see Section 5.5.3)
Changes in spatial sea level patterns through time may help
reconcile apparently inconsistent estimates of regional variations
in tide-gauge based sea level rise. For example, the minimum in
rise along the northwest Australian coast is consistent with the
results of Lambeck (2002) in having smaller rates of sea level
rise and indeed sea level fall off north-western Australia over
the last few decades. In addition, for the North Atlantic Ocean,
Figure 5.14. Variations in global mean sea level (difference to the mean 1993 to
mid-2001) computed from satellite altimetry from January 1993 to October 2005,
averaged over 65°S to 65°N. Dots are 10-day estimates (from the TOPEX/Posei-
don satellite in red and from the Jason satellite in green). The blue solid curve
corresponds to 60-day smoothing. Updated from Cazenave and Nerem (2004) and
Leuliette et al. (2004).
Observations: Oceanic Climate Change and Sea Level Chapter 5
the rate of rise reaches a maximum (over 2 mm yr–1) in a band
running east-northeast from the US east coast. The trends are
lower in the eastern than in the western Atlantic (Lambeck et
al., 1998; Woodworth et al., 1999; Mitrovica et al., 2001). Interannual and Decadal Variability and
Long-Term Changes in Sea Level
Sea level records contain a considerable amount of
interannual and decadal variability, the existence of which is
coherent throughout extended parts of the ocean. For example,
the global sea level curve in Figure 5.13 shows an approximately
10 mm rise and fall of global mean sea level accompanying the
1997–1998 ENSO event. Over the past few decades, the time
series of the rst EOF of Church et al. (2004) represents ENSO
variability, as shown by a signicant (negative) correlation with
the Southern Oscillation Index. The signature of the 1997–1998
El Niño is also clear in the altimetric maps of sea level anomalies
(see Section Model results suggest that large volcanic
eruptions produce interannual to decadal uctuations in the
global mean sea level (see Section 9.5.2).
Holgate and Woodworth (2004) concluded that the 1990s
had one of the fastest recorded rates of sea level rise averaged
along the global coastline (~4 mm yr–1), slightly higher than
the altimetry-based open ocean sea level rise (3 mm yr–1).
However, their analysis also shows that some previous decades
had comparably large rates of coastal sea level rise (e.g., around
1980; Figure 5.17). White et al. (2005) conrmed the larger sea
level rise during the 1990s around coastlines compared to the
open ocean but found that in some previous periods the coastal
rate was smaller than the open ocean rate, and concluded that
over the last 50 years the coastal and open ocean rates of change
were the same on average. The global reconstruction of Church
et al. (2004) and Church and White (2006) also exhibits large
decadal variability in the rate of global mean sea level rise,
and the 1993 to 2003 rate has been exceeded in some previous
decades (Figure 5.17). The variability is smaller in the global
reconstruction (standard deviation of overlapping 10-year rates
is 1.1 mm yr–1) than in the Holgate and Woodworth (2004)
coastal time series (standard deviation 1.7 mm yr–1). The rather
Figure 5.15. (a) Geographic distribution of short-term linear trends in mean
sea level (mm yr–1) for 1993 to 2003 based on TOPEX/Poseidon satellite altimetry
(updated from Cazenave and Nerem, 2004) and (b) geographic distribution of linear
trends in thermal expansion (mm yr–1) for 1993 to 2003 (based on temperature data
down to 700 m from Ishii et al., 2006).
Figure 5.16. (a) Geographic distribution of long-term linear trends in mean sea
level (mm yr–1) for 1955 to 2003 based on the past sea level reconstruction with tide
gauges and altimetry data (updated from Church et al., 2004) and (b) geographic
distribution of linear trends in thermal expansion (mm yr–1) for 1955 to 2003 (based
on temperature data down to 700 m from Ishii et al., 2006). Note that colours in (a)
denote 1.6 mm yr–1 higher values than those in (b).
Chapter 5 Observations: Oceanic Climate Change and Sea Level
low temporal correlation (r = 0.44) between the two time series
suggests that the statistical uncertainty in the linear trends
calculated from either data set probably underestimates the
systematic uncertainty in the results (Section 5.5.6).
Interannual or longer variability is a major reason why no
long-term acceleration of sea level has been identied using
20th-century data alone (Woodworth, 1990; Douglas, 1992).
Another possibility is that the sparse tide gauge network may
have been inadequate to detect it if present (Gregory et al.,
2001). The longest records available from Europe and North
America contain accelerations of the order of 0.4 mm yr–1 per
century between the 19th and 20th century (Ekman, 1988;
Woodworth et al., 1999). For the reconstruction shown in
Figure 5.13, Church and White (2006) found an acceleration
of 1.3 ± 0.5 mm yr–1 per century over the period 1870 to 2000.
These data support an inference that the onset of acceleration
occurred during the 19th century (see Section 9.5.2).
Geological observations indicate that during the last 2,000
years (i.e., before the recent rise recorded by tide gauges),
sea level change was small, with an average rate of only 0.0
to 0.2 mm yr–1 (see Section 6.4.3). The use of proxy sea level
data from archaeological sources is well established in the
Mediterranean. Oscillations in sea level from 2,000 to 100 yr
before present did not exceed ±0.25 m, based on the Roman-
Byzantine-Crusader well data (Sivan et al., 2004). Many
Roman and Greek constructions are relatable to the level of the
sea. Based on sea level data derived from Roman sh ponds,
which are considered to be a particularly reliable source of such
information, together with nearby tide gauge records, Lambeck
et al. (2004) concluded that the onset of the modern sea level
rise occurred between 1850 and 1950. Donnelly et al. (2004)
and Gehrels et al. (2004), employing geological data from
Connecticut, Maine and Nova Scotia salt-marshes together with
nearby tide gauge records, demonstrated that the sea level rise
observed during the 20th century was in excess of that averaged
over the previous several centuries.
The joint interpretation of the geological observations, the
longest instrumental records and the current rate of sea level
rise for the 20th century gives a clear indication that the rate of
sea level rise has increased between the mid-19th and the mid-
20th centuries. Regional Sea Level Change
Two regions are discussed here to give examples of local
variability in sea level: the northeast Atlantic and small Pacic
Interannual variability in northeast Atlantic sea level records
exhibits a clear relationship to the air pressure and wind changes
associated with the NAO, with the magnitude and sign of the
response depending primarily upon latitude (Andersson, 2002;
Wakelin et al., 2003; Woolf et al., 2003). The signal of the
NAO can also be observed to some extent in ocean temperature
records, suggesting a possible, smaller NAO inuence on
regional mean sea level via steric (density) changes (Tsimplis
et al., 2006). In the Russian Arctic Ocean, sea level time series
for recent decades also have pronounced decadal variability
that correlates with the NAO index. In this region, wind stress
and atmospheric pressure loading contribute nearly half of the
observed sea level rise of 1.85 mm yr–1 (Proshutinsky et al.,
Small Pacic Islands are the subject of much concern in
view of their vulnerability to sea level rise. The Pacic Ocean
region is the centre of the strongest interannual variability of the
climate system, the coupled ocean-atmosphere ENSO mode.
There are only a few Pacic Island sea level records extending
back to before 1950. Mitchell et al. (2001) calculated rates of
relative sea level rise for the stations in the Pacic region. Using
their results (from their Table 1) and focusing on only the island
stations with more than 50 years of data (only 4 locations), the
average rate of sea level rise (relative to the Earth’s crust) is
1.6 mm yr–1. For island stations with record lengths greater than
25 years (22 locations), the average rate of relative sea level rise
is 0.7 mm yr–1. However, these data sets contain a large range
of rates of relative sea level change, presumably as a result of
poorly quantied vertical land motions.
An example of the large interannual variability in sea level
is Kwajalein (8°44’N, 167°44’E) (Marshall Archipelago).
As shown in Figure 5.18, the local tide gauge data, the sea
level reconstructions of Church et al. (2004) and Church and
White (2006) and the shorter satellite altimeter record all
agree and indicate that interannual variations associated with
ENSO events are greater than 0.2 m. The Kwajalein data also
suggest increased variability in sea level after the mid-1970s,
consistent with the trend towards more frequent, persistent
and intense ENSO events since the mid-1970s (Folland et
al., 2001). For the Kwajalein record, the rate of sea level rise,
after correction for GIA land motions and isostatic response to
atmospheric pressure changes, is 1.9 ± 0.7 mm yr–1. However,
Figure 5.17. Overlapping 10-year rates of global sea level change from tide gauge
data sets (Holgate and Woodworth, 2004, in solid black; Church and White, 2006, in
dashed black) and satellite altimetry (updated from Cazenave and Nerem, 2004, in
green), and contributions to global sea level change from thermal expansion (Ishii et
al., 2006, in solid red; Antonov et al., 2005, in dashed red) and climate-driven land
water storage (Ngo-Duc et al., 2005, in blue). Each rate is plotted against the middle
of its 10-year period.
Observations: Oceanic Climate Change and Sea Level Chapter 5
the uncertainties in rates of sea level change increase rapidly
with decreasing record length and can be several mm yr–1
for decade-long records (depending on the magnitude of
the interannual variability). Sea level change on the atolls of
Tuvalu (western Pacic) has been the subject of intense interest
as a result of their low-lying nature and increasing incidence of
ooding. There are two records available at Funafuti, Tuvalu;
the rst record commences in 1977 and the second (with
rigorous datum control) in 1993. After allowing for subsidence
affecting the rst record, Church et al. (2006) estimate sea level
rise at Tuvalu to be 2.0 ± 1.7 mm yr–1, in agreement with the
reconstructed rate of sea level rise. Changes in Extreme Sea Level
Societal impacts of sea level change primarily occur via
the extreme levels rather than as a direct consequence of mean
sea level changes. Apart from non-climatic events such as
tsunamis, extreme sea levels occur mainly in the form of storm
surges generated by tropical or extra tropical cyclones. Secular
changes and decadal variability in storminess are discussed in
Chapter 3. Studies of variations in extreme sea levels during
the 20th century based on tide gauge data are fewer than studies
of changes in mean sea level for several reasons. A study on
changes in extremes, which are caused by changes in mean
sea level as well as changes in surges, is more complex than
the study of mean sea level changes. Moreover, the hourly
sampling interval normally used in tide gauge records is not
always sufcient to accurately capture the true extreme. Among
the different parameters often used to describe extremes, annual
maximum surge is a good indicator of climatic trends. For
study of long records extending back to the 19th century or
before, annual maximum surge-at-high-water (dened as the
maximum of the difference between observed high water and
the predicted tide at high water) is a better-suited parameter
because during that period high waters and not the full tidal
curve were recorded.
Studies of the longest records of extremes are inevitably
restricted to a small number of locations. From observed sea level
extremes at Liverpool since 1768, Woodworth and Blackman
(2002) concluded that the annual maximum surge-at-high-water
was larger in the late 18th, late 19th and late 20th centuries
than for most of the 20th century, qualitatively consistent with
the long-term variability in storminess from meteorological
data. From the tide gauge record at Brest from 1860 to 1994,
Bouligand and Pirazzoli (1999) found an increasing trend in
annual maxima and 99th percentile of surges; however, a
decreasing trend was found during the period 1953 to 1994.
From non-tidal residuals (‘surges’) at San Francisco since 1858,
Bromirski et al. (2003) concluded that extreme winter residuals
have exhibited a signicant increasing trend since about 1950,
a trend that is attributed to an increase in storminess during
this period. Zhang et al. (2000) concluded from records at 10
stations along the east coast of the USA since 1900 that the rise
in extreme sea level closely followed the rise in mean sea level.
A similar conclusion can be drawn from a recent study of Firing
and Merrield (2004), who found long-term increases in the
number and height of daily extremes at Honolulu (interestingly,
the highest-ever value being due an anticyclonic oceanic eddy
system in 2003), but no evidence for an increase relative to the
underlying upward mean sea level trend.
An analysis of 99th percentiles of hourly sea level at 141
stations over the globe for recent decades (Woodworth and
Blackman, 2004) showed that there is evidence for an increase
in extreme high sea level worldwide since 1975. In many cases,
the secular changes in extremes were found to be similar to
those in mean sea level. Likewise, interannual variability in
extremes was found to be correlated with regional mean sea
level, as well as to indices of regional climate patterns.
5.5.3 Ocean Density Changes
Sea level will rise if the ocean warms and fall if it cools,
since the density of the water column will change. If the
thermal expansivity were constant, global sea level change
would parallel the global ocean heat content discussed in
Section 5.2. However, since warm water expands more than
cold water (with the same input of heat), and water at higher
pressure expands more than at lower pressure, the global sea
level change depends on the three-dimensional distribution of
ocean temperature change.
Analysis of the last half century of temperature observations
indicates that the ocean has warmed in all basins (see Section
5.2). The average rate of thermosteric sea level rise caused
by heating of the global ocean is estimated to be 0.40 ±
0.09 mm yr–1 over 1955 to 1995 (Antonov et al., 2005), based
on ve-year mean temperature data down to 3,000 m. For the
0 to 700 m layer and the 1955 to 2003 period, the averaged
thermosteric trend, based on annual mean temperature data
from Levitus et al. (2005a), is 0.33 ± 0.07 mm yr–1 (Antonov
et al., 2005). For the same period and depth range, the mean
thermosteric rate based on monthly ocean temperature data
from Ishii et al. (2006) is 0.36 ± 0.12 mm yr–1. Figure 5.19
Figure 5.18. Monthly mean sea level curve for 1950 to 2000 at Kwajalein (8°44’N,
167°44’E). The observed sea level (from tide gauge measurements) is in blue, the
reconstructed sea level in red and the satellite altimetry record in green. Annual and
semi-annual signals have been removed from each time series and the tide gauge
data have been smoothed. The figure was drawn using techniques in Church et al.
(2004) and Church and White (2006).
Chapter 5 Observations: Oceanic Climate Change and Sea Level
shows the thermosteric sea level curve over 1955
to 2003 for both the Levitus and Ishii data sets.
The rate of thermosteric sea level rise is clearly not
constant in time and shows considerable uctuations
(Figure 5.17). A rise of more than 20 mm occurred
from the late 1960s to the late 1970s (giving peak
10-year rates in the early 1970s) with a smaller drop
afterwards. Another large rise began in the 1990s,
but after 2003, the steric sea level is decreasing
in both estimates (peak rates in the late 1990s).
Overlapping 10-year rates from these two estimates
have a very high temporal correlation (r = 0.97) and
the standard deviation of the rates is 0.7 mm yr–1.
The Levitus and Ishii data sets both give 0.32
± 0.09 mm yr–1 for the upper 700 m during 1961
to 2003, but the Levitus data set of temperature
down to 3,000 m ends in 1998. From the results of
Antonov et al. (2005) for thermal expansion, the
difference between the trends in the upper 3,000
m and the upper 700 m for 1961 to 1998 is about
0.1 mm yr–1. Assuming that the ocean below 700
m continues to contribute beyond 1998 at a similar
rate, with an uncertainty similar to that of the upper-
ocean contribution, we assess the thermal expansion
of the ocean down to 3,000 m during 1961 to 2003
as 0.42 ± 0.12 mm yr–1.
For the recent period 1993 to 2003, a value of 1.2 ±
0.5 mm yr–1 for thermal expansion in the upper 700 m is estimated
both by Antonov et al. (2005) and Ishii et al. (2006). Willis et
al. (2004) estimate thermal expansion to be 1.6 ± 0.5 mm yr–1,
based on combined in situ temperature proles down to 750 m
and satellite measurements of altimetric height. Including the
satellite data reduces the error caused by the inadequate sampling
of the prole data. Error bars were estimated to be about 2 mm
for individual years in the time series, with most of the remaining
error due to inadequate prole availability. A close result (1.8 ±
0.4 mm yr–1 steric sea level rise for 1993 to 2003) was recently
obtained by Lombard et al. (2006), based on a combined analysis
of in situ hydrographic data and satellite sea surface height and
SST data (Guinehut et al., 2004). It is presently unclear why the
latter two estimates are signicantly larger than the thermosteric
rates based on temperature data alone. It is possible that the in situ
data underestimate thermal expansion because of poor coverage
in Southern Oceans, and it is interesting to note that a model
based on assimilation of hydrographic data yields a somewhat
higher estimate of 2.3 mm yr–1 (Carton et al., 2005). Published
estimates of the steric sea level rates for 1955 to 2003 and 1993
to 2003 are shown in Table 5.2.
We assess the thermal expansion of the upper 700 m during
1993 to 2003 as 1.5 ± 0.5 mm yr–1, and that of the upper 3,000
m as 1.6 ± 0.5 mm yr–1, allowing for the ocean below 700 m as
for the earlier period (see also Section 5.5.6, Table 5.3).
Table 5.2. Recent estimates for steric sea level trends from different studies.
Steric sea level change
Reference with errors (mm yr–1) Period Depth range (m) Data Source
Antonov et al. (2005) 0.40 ± 0.09 1955–1998 0–3,000 Levitus et al. (2005b)
Antonov et al. (2005) 0.33 ± 0.07 1955–2003 0–700 Levitus et al. (2005b)
Ishii et al. (2006) 0.36 ± 0.06 1955–2003 0–700 Ishii et al. (2006)
Antonov et al. (2005) 1.2 ± 0.5 1993–2003 0–700 Levitus et al. (2005b)
Ishii et al. (2006) 1.2 ± 0.5 1993–2003 0–700 Ishii et al. (2006)
Willis et al. (2004) 1.6 ± 0.5 1993–2003 0–750 Willis et al. (2004)
Lombard et al. (2006) 1.8 ± 0.4 1993-2003 0-700 Guinehut et al. (2004)
Figure 5.19. Global sea level change due to thermal expansion for 1955 to 2003, based on Levi-
tus et al. (2005a; black line) and Ishii et al. (2006; red line) for the 0 to 700 m layer, and based on
Willis et al. (2004; green line) for the upper 750 m. The shaded area and the vertical red and green
error bars represent the 90% confidence interval. The black and red curves denote the deviation
from their 1961 to 1990 average, the shorter green curve the deviation from the average of the
black curve for the period 1993 to 2003.
Observations: Oceanic Climate Change and Sea Level Chapter 5
Antonov et al. (2002) attributed about 10% of
the global average steric sea level rise during recent
decades to halosteric expansion (i.e., the volume
increase caused by freshening of the water column). A
similar result was obtained by Ishii et al. (2006) who
estimated a halosteric contribution to 1955 to 2003
sea level rise of 0.04 ± 0.02 mm yr–1. While it is of
interest to quantify this effect, only about 1% of the
halosteric expansion contributes to the global sea level
rise budget. This is because the halosteric expansion
is nearly compensated by a decrease in volume of the
added freshwater when its salinity is raised (by mixing)
to the mean ocean value; the compensation would be
exact for a linear state equation (Gille, 2004; Lowe
and Gregory, 2006). Hence, for global sums of sea
level change, halosteric expansion cannot be counted
separately from the volume of added land freshwater
(which Antonov et al., 2002, also calculate; see
Section However, for regional changes in sea
level, thermosteric and halosteric contributions can be
comparably important (see, e.g., Section
5.5.4 Interpretation of Regional
Variations in the Rate of Sea
Level Change
Sea level observations show that whatever the time
span considered, rates of sea level change display
considerable regional variability (see Sections
and A number of processes can cause regional
sea level variations. Steric Sea Level Changes
Like the sea level trends observed by satellite
altimetry (see Section, the global distribution
of thermosteric sea level trends is not spatially
uniform. This is illustrated by Figure 5.15b and Figure 5.16b,
which show the geographical distribution of thermosteric sea
level trends over two different periods, 1993 to 2003 and 1955
to 2003 respectively (updated from Lombard et al., 2005). Some
regions experienced sea level rise while others experienced a
fall, often with rates that are several times the global mean.
However, the patterns of thermosteric sea level rise over the
approximately 50-year period are different from those seen in
the 1990s. This occurs because the spatial patterns, like the
global average, are also subject to decadal variability. In other
words, variability on different time scales may have different
characteristic patterns.
An EOF analysis of gridded thermosteric sea level time series
since 1955 (updated from Lombard et al., 2005) displays a spatial
pattern that is similar to the spatial distribution of thermosteric
sea level trends over the same time span (compare Figure 5.20
with Figure 5.16b). In addition, the rst principal component is
negatively correlated with the Southern Oscillation Index. Thus,
it appears that ENSO-related ocean variability accounts for the
largest fraction of variance in spatial patterns of thermosteric
sea level. Similarly, decadal thermosteric sea level in the North
Pacic and North Atlantic appears strongly inuenced by the
PDO and NAO respectively.
For the recent years (1993–2003), the geographic distribution
of observed sea level trends (Figure 5.15a) shows correlation
with the spatial patterns of thermosteric sea level change
(Figure 5.15b). This suggests that at least part of the non-
uniform pattern of sea level rise observed in the altimeter data
over the past decade can be attributed to changes in the ocean’s
thermal structure, which is itself driven by surface heating
effects and ocean circulation. Note that the steric changes due
to salinity changes have not been included in these gures due
to insufcient salinity data in parts of the World Ocean.
Ocean salinity changes, while unimportant for sea level at
the global scale, can have an effect on regional sea level (e.g.,
Antonov et al., 2002; Ishii et al., 2006; Section 5.5.3). For
example, in the subpolar gyre of the North Atlantic, especially
in the Labrador Sea, the halosteric contribution nearly
Figure 5.20. (a) First mode of the EOF decomposition of the gridded thermosteric sea level
time series of yearly temperature data down to 700 m from Ishii et al. (2006). (b) The normalised
principal component (black solid curve) is highly correlated with the negative Southern Oscillation
Index (dotted red curve).
Chapter 5 Observations: Oceanic Climate Change and Sea Level
counteracts the thermosteric contribution. This observational
result is supported by results from data assimilation into models
(e.g., Stammer et al., 2003). Since density changes can result
not only from surface buoyancy uxes but also from the wind,
a simple attribution of density changes to buoyancy forcing is
not possible.
While much of the non-uniform pattern of sea level change
can be attributed to thermosteric volume changes, the difference
between observed and thermosteric spatial trends show a high
residual signal in a number of regions, especially in the southern
oceans. Part of these residuals is likely due to the lack of ocean
temperature coverage in remote oceans as well as in deep layers
(below 700 m), and to regional salinity change. Ocean Circulation Changes
The highly non-uniform geographical distribution of steric
sea level trends is closely connected, through geostrophic
balance, with changes in ocean surface circulation. Density and
circulation changes result from changes in atmospheric forcing
that is primarily by surface wind stress and buoyancy ux
(i.e., heat and freshwater uxes). The wind alone can therefore
cause local (but not global) changes in steric sea level. Ocean
general circulation models based on the assimilation of ocean
data satisfactorily reproduce the spatial structure of sea level
trends for the past decade, and show in particular that the
tropical Pacic pattern results from decadal uctuations in the
depth of the tropical thermocline and change in equatorial trade
winds (Carton et al., 2005; Köhl et al., 2006). The similarity of
the patterns of steric and actual sea level change indicates that
density changes are the dominant inuence. Discrepancies may
indicate a signicant contribution from changes in the wind-
driven barotropic circulation, especially at high latitudes. Surface Atmospheric Pressure Changes
Surface atmospheric pressure also causes regional sea level
variations. Over time scales longer than a few days, the ocean
adjusts nearly isostatically to changes in atmospheric pressure
(inverted barometer effect), that is, for each 1 hPa sea level
pressure increase the ocean is depressed by approximately 10
mm, shifting the underlying mass sideways to other regions.
For the temporal average, regional changes in sea level caused
by atmospheric pressure loading reach about 0.2 m (e.g.,
between the subtropical Atlantic and the subpolar Atlantic).
Such effects are generally corrected for in tide gauge and
altimetry-based sea level analyses. The inverted barometer
effect has a negligible effect on global mean sea level, because
water is nearly incompressible, but is signicant when averaged
over the area of T/P and Jason-1 altimetry, which does not
cover the whole World Ocean (Ponte, 2006). For that reason,
the altimetry-based mean sea level curve is corrected for the
inverted barometer effect. Solid Earth and Geoid Changes
Geodynamical processes related to the solid Earth’s elastic
and viscoelastic response to spatially variable ice melt loading
(due to the last deglaciation and present-day land ice melt)
also cause non-uniform sea level change (e.g., Mitrovica et
al., 2001; Peltier, 2001, 2004; Plag, 2006). The solid Earth and
oceans continue to respond to the ice and complementary water
loads associated with the late Pleistocene and early Holocene
glacial cycles through GIA. This process not only drives large
crustal uplift near the location of former ice complexes, but also
produces a worldwide signature in sea level that results from
gravitational, deformational and rotational effects: as the viscous
mantle material ows to restore isostasy during and after the
last deglaciation, uplift occurs under the former centres of the
ice sheets while the surrounding peripheral bulges experience a
subsidence. The return of the melt water to the oceans produces
an ongoing geoid change resulting in subsidence of the ocean
basins and an upward warping of the continents, while the
ow of water into the subsiding peripheral bulges contributes
a broad scale sea level fall in the far eld of the ice complexes.
The combined gravitational and deformational effects also
perturb the rotation vector of the planet, and this perturbation
feeds back into variations in the position of the crust and the
geoid (an equipotential surface of the Earth’s gravity eld that
coincides with the mean surface of the oceans). Corrections for
GIA effects are made to both tide gauge and altimeter estimates
of global sea level change (see Sections and
Self-gravitation and deformation of the Earth’s surface in
response to the ongoing change in loading by glaciers and ice
sheets is another cause of regional sea level variations. Model
predictions show quite different patterns of non-uniform sea
level change depending on the source of the ice melt (Mitrovica
et al., 2001; Plag, 2006), and associated regional sea level
variations reach up to a few 0.1 mm yr–1.
5.5.5 Ocean Mass Change
Global mean sea level will rise if water is added to the ocean
from other reservoirs in the climate system. Water storage in
the atmosphere is equivalent to only about 35 mm of global
mean sea level, and the observed atmospheric storage trend
of about 0.04 mm yr–1 in recent decades (Section is
unimportant compared with changes in ice and water stored
on land, described in this subsection. Variations in land water
storage result from variations in climatic conditions, direct
human intervention in the water cycle and human modication
of the land surface. Ocean Mass Change Estimated from Salinity
Global salinity changes can be caused by changes in the
global sea ice volume (which do not inuence sea level) and by
ocean mass changes (which do). Thus in principle, global salinity
Observations: Oceanic Climate Change and Sea Level Chapter 5
changes can be used to estimate the global average sea level
change due to fresh water input (Antonov et al., 2002; Munk,
2003; Wadhams and Munk, 2004). However, the accuracy of
these estimates depends on the accuracy of the estimates for
both sea ice volume (Hilmer and Lemke, 2000; Wadhams and
Munk, 2004; see also Section 4.4) and global salinity change
(Section 5.2.3). We assess that the error in estimates of ocean
mass changes derived from salinity changes and sea ice melt is
too large to provide useful constraints on the sea level change
budget (Section 5.5.6). Land Ice
During the 20th century, glaciers and ice caps have
experienced considerable mass losses, with strong retreats in
response to global warming after 1970. For 1961 to 2003, their
contribution to sea level is assessed as 0.50 ± 0.18 mm yr–1 and
for 1993 to 2003 as 0.77 ± 0.22 mm yr–1 (see Section 4.5.2).
As discussed in Section and Table 4.6, the Greenland
Ice Sheet has also been losing mass in recent years, contributing
0.05 ± 0.12 mm yr–1 to sea level rise during 1961 to 2003 and
0.21 ± 0.07 mm yr–1 during 1993 to 2003. Assessments of
contributions to sea level rise from the Antarctic Ice Sheet are less
certain, especially before the advent of satellite measurements,
and are 0.14 ± 0.41 mm yr–1 for 1961 to 2003 and 0.21 ±
0.35 mm yr–1 for 1993 to 2003. Geodetic data on Earth rotation
and polar wander allow a late-20th century sea level contribution
of up to about 1 mm yr–1 from land ice (Mitrovica et al., 2006).
However, recent estimates of ice sheet mass change exclude
the large contribution inferred for Greenland by Mitrovica et
al. (2001) from the geographical pattern of sea level change,
conrming the lower rates reported above. Climate-Driven Change in Land Water Storage
Continental water storage includes water (both liquid
and solid) stored in subsurface saturated (groundwater) and
unsaturated (soil water) zones, in the snowpack, and in surface
water bodies (lakes, articial reservoirs, rivers, oodplains and
wetlands). Changes in concentrated stores, most notably very
large lakes, are relatively well known from direct observation.
In contrast, global estimates of changes in distributed surface
stores (soil water, groundwater, snowpack and small areas of
surface water) rely on computations with detailed hydrological
models coupled to global ocean-atmosphere circulation models
or forced by observations. Such models estimate the variation in
land water storage by solving the water balance equation. The
Land Dynamics (LaD) model developed by Milly and Shmakin
(2002) provides global by 1° monthly gridded time series
of root zone soil water, groundwater and snowpack for the last
two decades. With these data, the contributions of time-varying
land water storage to sea level rise in response to climate
change have been estimated, resulting in a small positive sea
level trend of about 0.12 mm yr–1 for the last two decades, with
larger interannual and decadal uctuations (Milly et al., 2003).
From a land surface model forced by a global climatic data set
based on standard reanalysis products and on observations,
land water changes during the past ve decades were found
to have low-frequency (decadal) variability of about 2 mm
in amplitude but no signicant trend (Ngo-Duc et al., 2005).
These decadal variations are related to groundwater and are
caused by precipitation variations. They are strongly negatively
correlated with the de-trended thermosteric sea level (Figure
5.17). This suggests that the land water contribution to sea level
and thermal expansion partly compensate each other on decadal
time scales. However, this conclusion depends on the accuracy
of the precipitation in reanalysis products. Anthropogenic Change in Land Water Storage
The amount of anthropogenic change in land water storage
systems cannot be estimated with much condence, as already
discussed by Church et al. (2001). A number of factors can
contribute to sea level rise. First, natural groundwater systems
typically are in a condition of dynamic equilibrium where, over
long time periods, recharge and discharge are in balance. When
the rate of groundwater pumping greatly exceeds the rate of
recharge, as is often the case in arid or even semi-arid regions,
water is removed permanently from storage. The water that is
lost from groundwater storage eventually reaches the ocean
through the atmosphere or surface ow, resulting in sea level
rise. Second, wetlands contain standing water, soil moisture and
water in plants equivalent to water roughly 1 m deep. Hence,
wetland destruction contributes to sea level rise. Over time
scales shorter than a few years, diversion of surface waters
for irrigation in the internally draining basins of arid regions
results in increased evaporation. The water lost from the basin
hydrologic system eventually reaches the ocean. Third, forests
store water in living tissue both above and below ground. When
a forest is removed, transpiration is eliminated so that runoff is
favoured in the hydrologic budget.
On the other hand, impoundment of water behind dams
removes water from the ocean and lowers sea level. Dams have
led to a sea level drop over the past few decades of –0.5 to –0.7
mm yr–1 (Chao, 1994; Sahagian et al., 1994). Inltration from
dams and irrigation may raise the water table, storing more
water. Gornitz (2001) estimated –0.33 to –0.27 mm yr–1 sea
level change equivalent held by dams (not counting additional
potential storage due to subsurface inltration).
It is very difcult to provide accurate estimates of the net
anthropogenic contribution, given the lack of worldwide
information on each factor, although the effect caused by
dams is possibly better known than other effects. According to
Sahagian (2000), the sum of the above effects could be of the
order of 0.05 mm yr–1 sea level rise over the past 50 years, with
an uncertainty several times as large.
In summary, our assessment of the land hydrology
contribution to sea level change has not led to a reduction in
the uncertainty compared to the TAR, which estimated the
rather wide ranges of –1.1 to +0.4 mm yr–1 for 1910 to 1990
Chapter 5 Observations: Oceanic Climate Change and Sea Level
and –1.9 to +1.0 mm yr–1 for 1990. However, indirect evidence
from considering other contributions to the sea level budget
(see Section 5.5.6) suggests that the land contribution either is
small (<0.5 mm yr–1) or is compensated for by unaccounted or
underestimated contributions.
5.5.6 Total Budget of the Global Mean Sea
Level Change
The various contributions to the budget of sea level change
are summarised in Table 5.3 and Figure 5.21 for 1961 to 2003
and 1993 to 2003. Some terms known to be small have been
omitted, including changes in atmospheric water vapour and
climate-driven change in land water storage (Section 5.5.5),
permafrost and sedimentation (see, e.g., Church et al., 2001),
which very likely total less than 0.2 mm yr–1. The poorly known
anthropogenic contribution from terrestrial water storage (see
Section is also omitted.
For 1961 to 2003, thermal expansion accounts for only
23 ± 9% of the observed rate of sea level rise. Miller and
Douglas (2004) reached a similar conclusion by computing
steric sea level change over the past 50 years in three oceanic
regions (northeast Pacic, northeast Atlantic and western
Atlantic); they found it to be too small by about a factor of
three to account for the observed sea level rise based on nine
tide gauges in these regions. They concluded that sea level rise
in the second half of the 20th century was mostly due to water
mass added to the oceans. However, Table 5.3 shows that the
sum of thermal expansion and contributions from land ice is
smaller by 0.7 ± 0.7 mm yr–1 than the observed global average
sea level rise. This is likely to be a signicant difference. The
assessment of Church et al. (2001) could allow this difference
to be explained by positive anthropogenic terms (especially
groundwater mining) but these are expected to have been
outweighed by negative terms (especially impoundment). We
conclude that the budget has not yet been closed satisfactorily.
Given the large temporal variability in the rate of sea level rise
evaluated from tide gauges (Section and Figure 5.17),
the budget is rather problematic on decadal time scales. The
thermosteric contribution has smaller variability (though still
substantial; Section 5.5.3) and there is only moderate temporal
correlation between the thermosteric rate and the tide gauge
rate. The difference between them has to be explained by ocean
mass change. Because the thermosteric and climate-driven land
water contributions are negatively correlated (Section,
Table 5.3. Estimates of the various contributions to the budget of global mean sea level change for 1961 to 2003 and 1993 to 2003 compared with the observed rate of rise.
Ice sheet mass loss of 100 Gt yr–1 is equivalent to 0.28 mm yr–1 of sea level rise. A GIA correction has been applied to observations from tide gauges and altimetry. For the sum,
the error has been calculated as the square root of the sum of squared errors of the contributions. The thermosteric sea level changes are for the 0 to 3,000 m layer of
the ocean.
Sea Level Rise (mm yr–1)
Source 1961–2003 1993–2003 Reference
Thermal Expansion 0.42 ± 0.12 1.6 ± 0.5 Section 5.5.3
Glaciers and Ice Caps 0.50 ± 0.18 0.77 ± 0.22 Section 4.5
Greenland Ice Sheet 0.05 ± 0.12 0.21 ± 0.07 Section 4.6.2
Antarctic Ice Sheet 0.14 ± 0.41 0.21 ± 0.35 Section 4.6.2
Sum 1.1 ± 0.5 2.8 ± 0.7
Observed 1.8 ± 0.5 Section
3.1 ± 0.7 Section
Difference (Observed –Sum) 0.7 ± 0.7 0.3 ± 1.0
Figure 5.21. Estimates of the various contributions to the budget of the global mean
sea level change (upper four entries), the sum of these contributions and the observed
rate of rise (middle two), and the observed rate minus the sum of contributions
(lower), all for 1961 to 2003 (blue) and 1993 to 2003 (brown). The bars represent the
90% error range. For the sum, the error has been calculated as the square root of the
sum of squared errors of the contributions. Likewise the errors of the sum and the
observed rate have been combined to obtain the error for the difference.
Observations: Oceanic Climate Change and Sea Level Chapter 5
the apparent difference implies contributions during some
10-year periods from land ice, the only remaining term,
exceeding 2 mm yr–1 (Figure 5.17). Since it is unlikely that
the land ice contributions of 1993 to 2003 were exceeded
in earlier decades (Figure 4.14 and Section, we
conclude that the maximum 10-year rates of global sea level
rise are likely overestimated from tide gauges, indicating
that the estimated variability is excessive.
For 1993 to 2003, thermal expansion is much larger and
land ice contributes 1.2 ± 0.4 mm yr–1. These increases
may partly reect decadal variability rather than an
acceleration (Section 5.5.3; attribution of changes in rates
and comparison with model results are discussed in Section
9.5.2). The sum is still less than the observed trend but the
discrepancy of 0.3 ± 1.0 mm yr–1 is consistent with zero. It is
interesting to note that the difference between the observed
total and thermal expansion (assumed to be due to ocean
mass change) is about the same in the two periods. The
more satisfactory assessment for recent years, during which
individual terms are better known and satellite altimetry is
available, indicates progress since the TAR.
5.6 Synthesis
The patterns of observed changes in global heat content and
salinity, sea level, steric sea level, water mass evolution and
biogeochemical cycles described in the previous four sections
are broadly consistent with known characteristics of the large-
scale ocean circulation (e.g., ENSO, NAO and SAM).
There is compelling evidence that the heat content of the
World Ocean has increased since 1955 (Section 5.2). In the
North Atlantic, the warming is penetrating deeper than in the
Pacic, Indian and Southern Oceans (Figure 5.3), consistent
with the strong convection, subduction and deep overturning
circulation cell that occurs in the North Atlantic Ocean. The
overturning cell in the North Atlantic region (carrying heat
and water downwards through the water column) also suggests
that there should be a higher anthropogenic carbon content as
observed (Figure 5.11). Subduction of SAMW (and to a lesser
extent AAIW) also carries anthropogenic carbon into the ocean,
which is observed to be higher in the formation areas of these
subantarctic water masses (Figure 5.10). The transfer of heat into
the ocean also leads to sea level rise through thermal expansion,
and the geographical pattern of sea level change since 1955 is
largely consistent with thermal expansion and with the change
in heat content (Figure 5.2).
Although salinity measurements are relatively sparse
compared with temperature measurements, the salinity data
also show signicant changes. In global analyses, the waters
at high latitudes (poleward of 50°N and 70°S) are fresher in
the upper 500 m (Figure 5.5 World). In the upper 500 m, the
subtropical latitudes in both hemispheres are characterised by
an increase in salinity. The regional analyses of salinity also
show a similar distributional change with a freshening of key
high-latitude water masses such as LSW, AAIW and NPIW, and
increased salinity in some of the subtropical gyres such as that
at 24°N. The North Atlantic (and other key ocean water masses)
also shows signicant decadal variations, such as the recent
increase in surface salinity in the North Atlantic subpolar gyre.
At high latitudes (particularly in the NH), there is an observed
increase in melting of perennial sea ice, precipitation, and
glacial melt water (see Chapter 4), all of which act to freshen
high-latitude surface waters. At mid-latitudes it is likely that
evaporation minus precipitation has increased (i.e., the transport
of freshwater from the ocean to the atmosphere has increased).
The pattern of salinity change suggests an intensication in the
Earth’s hydrological cycle over the last 50 years. These trends
are consistent with changes in precipitation and inferred greater
water transport in the atmosphere from low latitudes to high
latitudes and from the Atlantic to the Pacic.
Figure 5.22 shows zonal means of changes in temperature,
anthropogenic carbon, sea level rise and a passive tracer (CFC).
It is remarkable that these independent variables (albeit with
widely varying reference periods) show a common pattern
of change in the ocean. Specically, the close similarity of
higher levels of warming, sea level rise, anthropogenic carbon
and CFC-11 at mid-latitudes and near the equator strongly
suggests that these changes are the result of changes in ocean
ventilation and circulation. Warming of the upper ocean should
lead to a decrease in ocean ventilation and subduction rates, for
which there is some evidence from observed decreases in O2
In the equatorial Pacic, the pattern of steric sea level rise
also shows that strong west to east gradients in the Pacic
have weakened (i.e., it is now cooler in the western Pacic and
warmer in the eastern Pacic). This decrease in the equatorial
temperature gradient is consistent with a tendency towards
more prolonged and stronger El Niños over this same period
(see Section 3.6.2).
Figure 5.22. Averages of temperature change (blue, from Levitus et al., 2005a),
anthropogenic carbon (red, from Sabine et al., 2004b) and CFC-11 (green, from Willey et
al., 2004) along lines of constant latitude over the top 700-m layer of the upper ocean.
Also shown is sea level change averaged along lines of constant latitude (black, from
Cazenave and Nerem, 2004). The temperature changes are for the 1955 to 2003 period,
the anthropogenic carbon is since pre-industrial times (i.e., 1750), CFC-11 concentrations
are for the period 1930 to 1994 and sea level for the period 1993 to 2003.
Chapter 5 Observations: Oceanic Climate Change and Sea Level
The subduction of carbon into the ocean has resulted in
calcite and aragonite saturation horizons generally becoming
shallower and pH decreasing primarily in the surface and near-
surface ocean causing the ocean to become more acidic.
Since the TAR, the capability to measure most of the
processes that contribute to sea level has been developed. In
the 1990s, the observed sea level rise that was not explained
through steric sea level rise could largely be explained by the
transfer of mass from glaciers, ice sheets and river runoff (see
Section 5.5). Figure 5.23 is a schematic that summarises the
observed changes.
All of these observations taken together give high condence
that the ocean state has changed, that the spatial distribution of
the changes is consistent with the large-scale ocean circulation
and that these changes are in response to changed ocean surface
While there are many robust ndings regarding the changed
ocean state, key uncertainties still remain. Limitations in ocean
sampling (particularly in the SH) mean that decadal variations in
global heat content, regional salinity patterns, and rates of global
sea level rise can only be evaluated with moderate condence.
Furthermore, there is low condence in the evidence for trends
in the MOC and the global ocean freshwater budget. Finally,
the global average sea level rise for the last 50 years is likely to
be larger than can be explained by thermal expansion and loss
of land ice due to increased melting, and thus for this period it
is not possible to satisfactorily quantify the known processes
causing sea level rise.
Figure 5.23. Schematic of the observed changes in the ocean state, including ocean temperature, ocean salinity, sea level, sea ice and biogeochemical cycles. The legend
identifies the direction of the changes in these variables.
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