Originally published by our Content Partner: Intergovernmental Panel on Climate Change (other articles)

Observations: Oceanic Climate Change and Sea Level

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.

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 salinification of shallower parts of the tropical and subtropical oceans. Freshening is pronounced in the Pacific 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 Pacific. 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 Pacific has been observed. Long-term cooling is observed in the North Atlantic subpolar gyre and in the central North Pacific. 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 significantly 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 firm 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 confidence 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, significantly 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 significantly 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.

• 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.

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) influences 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 influenced 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-stratified 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 Pacific 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 density.

The Third Assessment Report (TAR) discussed some aspects of the ocean’s role. Folland et al. (2001) concluded that the global ocean has significantly 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 significant 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% confidence 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 defined 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., 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 profiling floats, 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

5.2.2.1 Long-Term Temperature Changes

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 representing 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 error 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.

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 profiles 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 confidence 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 significant increasing trend in ocean heat content.

The data used in estimating the Levitus et al. (2005a) ocean temperature fields (for the above heat content estimates) do not include sea surface temperature (SST) observations, 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 confidence 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 five-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 surface.

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.

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 Pacific sectors has generally warmed. The Pacific Ocean is characterised by warming with major exceptions along 40°N and the western tropical Pacific.

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).

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 fields (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 Pacific centred on the equator and 150 m depth. Cooling also occured in the 32°N to 48°N region of the Pacific Ocean and the 49°N to 60°N region of the Atlantic Ocean. Regional temperature changes are discussed further in Section 5.3.

5.2.2.2 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 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 Pacific 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 Pacific 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 confidence that there is substantial inter-decadal variability in global ocean heat content.

5.2.2.3 Implications for Earth’s Heat Balance

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 energy 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 results 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).

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 (contributing about 40%), the energy associated with ice melt contributes only about 1% to the Earth’s energy budget.

5.2.3 Ocean Salinity

Figure 5.5. Linear trends (1955–1998) of zonally averaged salinity (psu) in the upper 500 m of the Atlantic, Pacifi c, 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).

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 fields) 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 profiles 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 Pacific with the exception of the South Pacific 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 Pacific subtropical region is becoming saltier, on average the whole water column in the Pacific Basin is becoming fresher (Boyer et al., 2005). The increasing difference in volume-averaged salinity between the Atlantic and Pacific Oceans suggests changes in freshwater transport between these two ocean basins.

We are confident 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 insufficient 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 Pacific (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 flux (see Section 5.2.2.1). At regional scales, few estimates of heat flux changes have been possible. During the last 50 years, net heat fluxes from the ocean to the atmosphere demonstrate locally decreasing values (up to 1 W m^–2 yr^–1) over the southern flank 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 flux observations is insufficient to permit a direct assessment of changes in heat flux. Air-sea fluxes are discussed in Section 3.5.6.

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 flux fields (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

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 differences 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 masse shown include Antarctic Intermediate Water (AAIW), Circumpolar Deep Water (CDW), North Atlantic Deep Water (NADW), Mediterranean Water (MW), Labrador Sea Water (LSW), Denmark Strait Overflow Water (DSOW) and North Pacifi c Intermediate Water NPIW). The areas shaded in grey represent the seafloor and oceanic crust.

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 Pacific (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 fill 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 Pacific section show widespread change, but the North Pacific 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 the ocean’s vertical stratification 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 Pacific 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 Pacific Oceans. In the Northern Hemisphere (NH), the Pacific intermediate waters have freshened, and the underlying deep waters did not change, consistent with no local bottom water source in the North Pacific. In the central North Atlantic, the intermediate layer (approximately 900–1,200 m) became saltier due to increased salinity in the outflow from the Mediterranean that feeds this layer.

5.3.2 Atlantic and Arctic Oceans

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 identified from the global analyses of SST (see Section 3.2.2.3). 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).

5.3.2.1 North Atlantic Subpolar Gyre, Labrador Sea and Nordic Seas

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.

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 influence 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 significant 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).

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 overflow 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 intensifi cation of the subpolar gyre. Since 1994, the salinity of the inflow 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 overflows from the upper 1,500 m of the Nordic Seas through the Denmark Strait and Faroe Channel. The marked freshening of the overflow 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 overflow 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 overflows 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 (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).

5.3.2.2 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 Pacific 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 stratification 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 significant decadal variability in the Arctic Ocean, no systematic long-term trend in subsurface arctic waters has been identified.

5.3.2.3 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 reflect 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 mid-1990s.

Intermediate water (800–1,200 m) in the mid-latitude eastern North Atlantic is strongly influenced by the saline Mediterranean Water (MW; Section 5.3.2.3). This saline layer joins the southward-flowing 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 5.3.5.2), 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 influenced by a South Atlantic dipole in SST (Venegas et al., 1998), associated with latent heat fluxes 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 longterm trends in temperature and salinity (Section 5.2), decadal variations are an important climate signal for this region.

5.3.2.4 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 reflected 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 inflow (Rohling and Bryden, 1992). These changes in water properties and circulation are linked to the long-term variability of surface fluxes (Krahmann and Schott, 1998) with contributions from the NAO (Vignudelli et al., 1999) that produce consistent changes in surface heat fluxes 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 outflow of water into the North Atlantic at Gibraltar (see also Section 5.3.2.3). Part of this shift in Mediterranean outflow 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 flux 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 outflow 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 Pacific 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 Pacific 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 Pacific has cooled along 40°N. Long-term trends are rather difficult to discern in the upper Pacific 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 Pacific is freshening but there are embedded salinity increases in the subtropical upper ocean, where strong evaporation dominates.

5.3.3.1 Pacific Upper Ocean Changes

In the North Pacific, 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 Pacific 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 Pacific Ocean, a positive PDO state such as occurred after 1976 is characterised by a strengthened Kuroshio Extension. After 1976, the Kuroshio Extension and North Pacific 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 (North Pacific 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 Pacific, 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 Pacific 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 Pacific 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 _mol kg^–1 yr^–1 (Gamo et al., 1986; Minami et al., 1998). However, possibly because of weakened vertical stratification at mid-depths associated with the decades-long warming, deepwater 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 deepwater production in subsequent years.

5.3.3.2 Intermediate and Deep Circulation and Water Property Changes

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).

Since the 1970s, the major mid-depth water mass in the North Pacific, North Pacific 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 Pacific, with most influence from the Okhotsk Sea, and reflects changes in northern North Pacific 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 Pacific, 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 Pacific 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 largescale, significant warming of the bottom 1,000 m across the entire North Pacific 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 Pacific 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 Pacific (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 significant 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 significantly 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 Pacific passing through the Indonesian Seas to the Indian Ocean and then onward into the South Atlantic. Much of this throughflow occurs in the tropics south of the equator, and is strongly affected by ENSO and the Indian Ocean Dipole (see Section 3.6.7.2). 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-flowing throughflow water.

Long-term trends in transport and properties of the throughflow 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 Pacific 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 across the equator is accomplished by a wind-driven, shallow cross-equatorial cell; data assimilation analysis has shown a significant 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 Pacific 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 Pacific Oceans are related to changes in the Southern Ocean waters but have largely been described in those sections.

5.3.5.1 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 Pacific 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 stratification 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 float profile 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).

5.3.5.2 Antarctic Regions and Antarctic Circumpolar Current

The ACC, the longest current system in the world, has a transport through Drake Passage of about 130 Sv, with significant 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 (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 5.3.2.3). 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, significant 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 significantly 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

5.3.6.1 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 reflected 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 confidence 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 Pacific 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 Pacific 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 Pacific subpolar gyre north of 45°N is apparent in both regional analyses (Section 5.3.3.2) 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 fields 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 difficult 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 Pacific, freshening at high latitudes and increased evaporation in the subtropics, cooling in the central North Pacific, warming in the eastern and tropical Pacific 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 influence normally associated with the NAO, PDO and SAM, suggests that these coherent changes between both hemispheres are associated with a global phenomenon.

5.3.6.2 Consistency with the Large-Scale Ocean Circulation

The observed changes are broadly consistent with scientific 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 significant 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.

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 acidification is poorly known and could cause further changes in the marine carbon cycle with consequences that are difficult 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 fluxes of O₂ indirectly based on heat flux 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 fluxes of N₂O and CH₄ (including CH₄ hydrates) are discussed in Chapter 7.

5.4.2 Carbon

5.4.2.1 Total Change in Dissolved Inorganic Carbon and Air-Sea Carbon Dioxide Flux

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.

Direct observations of oceanic dissolved inorganic carbon (DIC; i.e., the sum of CO₂ plus carbonate and bicarbonate) reflect 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 7.3.2.4 for quantitative estimates). In the equatorial Pacific, 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 five 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 flux of 1.6 ± 1 GtC yr^–1 for the year 1995 (Takahashi et al., 2002; Section 7.3.2.3.2, Figure 7.8). It is not yet possible to detect large-scale changes in the global air-sea CO₂ flux from direct observations because of the large influence of climate variability. However, estimates from inverse methods of the air-sea CO₂ flux from the spatio-temporal distribution of atmospheric CO₂ suggest that the global air-sea CO₂ flux 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).

5.4.2.2 Anthropogenic Carbon Change

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.

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 reflect 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 Pacific (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).

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 profiles 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 identified, mostly caused by assumptions about the time evolution of CO₂, the age or the identification of water masses (Matsumoto and Gruber, 2005), and the recent changes in surface warming and stratification (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 identified.

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).

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 chlorofluorocarbons (CFCs) of atmospheric origin (Willey et al., 2004), confirming 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 include all emissions that have an influence 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