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Diagnosing Natural Variability of North Atlantic Water Masses in HadCM3 KEITH HAINES
AND
CHRIS OLD
Environmental Systems Science Centre, The University of Reading, Reading, United Kingdom (Manuscript received 9 December 2003, in final form 26 October 2004) ABSTRACT A study of thermally driven water mass transformations over 100 yr in the ocean component of the Third Hadley Centre Coupled Ocean–Atmosphere General Circulation Model (HadCM3) is presented. The processes of surface-forced transformations, subduction and mixing, both above and below the winter mixed layer base, are quantified. Subtropical Mode Waters are formed by surface heat fluxes and subducted at more or less the same rate. However, Labrador Seawater and Nordic Seawater classes (the other main subduction classes) are primarily formed by mixing within the mixed layer with very little formation directly from surface heat fluxes. The Subpolar Mode Water classes are dominated by net obduction of water back into the mixed layer from below. Subtropical Mode Water (18°C) variability shows a cycle of formation by surface fluxes, subduction ⬃2 yr later, followed by mixing with warmer waters below the winter mixed layer base during the next 3 yr, and finally obduction back into the mixed layer at 21°C, ⬃5 yr after the original formation. Surface transformation of Subpolar Mode Waters, ⬃12°C, are led by surface transformations of warmer waters by up to 5 yr as water is transferred from the subtropical gyre. They are also led by obduction variability from below the mixed layer, by ⬃2 yr. The variability of obduction in Subpolar Mode Waters also appears to be preceded, by 3–5 yr, by variability in subduction of Labrador Sea Waters at ⬃6°C. This supports a mechanism in which southward-propagating Labrador seawater anomalies below the subpolar gyre can influence the upper water circulation and obduction into the mixed layer.
1. Introduction The Lagrangian conserved properties of ocean water masses within the main thermocline vary only on long time scales, and therefore a record of these water mass properties and volumes contain a record of air–sea forcing and interior mixing processes going on in the ocean. It is a challenge to unravel this water mass record to learn as much as possible of the changes in ocean processes. The long-term aim is thus to fully utilize the record of ocean properties available from historical hydrography and the results of ocean reanalysis projects based on these observations. If heat content anomalies are isolated from the atmosphere and do not disperse, then when they again contact the atmosphere the stored heat can be released, giving a potentially predictable atmospheric response. Many papers have focused on ocean temperature and heat content anomalies, and the hydrographic record Corresponding author address: Keith Haines, Environmental Systems Science Centre, The University of Reading, P.O. Box 238, Reading RG6 6AL, United Kingdom. E-mail:
[email protected]
© 2005 American Meteorological Society
JCLI3348
shows large-scale temperature anomalies developing and propagating in the North Atlantic (Deser and Blackmon 1993; Kushnir 1994; Sutton and Allen 1997; Grey et al. 2000; Hakkinen 2000) and in the North Pacific (Schneider et al. 1999) on interannual to decadal time scales. Modeling studies (Krahmann et al. 2001; Dong and Sutton 2002; Cooper and Gordon 2002) have also shown heat anomalies in the upper 400 m of the North Atlantic, formed in response to large-scale atmospheric variability, persisting and propagating along the Gulf Stream and North Atlantic Current. However, anomalies in the volume and properties of water masses, broken down into temperature classes, represent a more complete record of ocean thermodynamic behavior, as well as implicitly containing the heat content information, and we shall focus on a water mass description in this paper. Most work on water mass volumes focus on “mode” waters, which appear as modes in water census diagnostics (e.g., Worthington 1981). They form layers of weak stratification and homogeneous properties within the thermocline of the ocean (Hanawa and Talley 2001). Most mode waters are formed at the surface in
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regions with a large exchange of heat between the ocean and atmosphere during winter, and they enter the thermocline by advection (subduction) across the winter mixed layer base by Ekman pumping and the geostrophic currents of the wind-driven gyres (e.g., Stommel 1979; Woods 1985; Williams 1991; Williams et al. 1995; Marshall 1997). Their weak stratification and (hence low potential vorticity), is preserved as they move around the gyre. The predominance of mode waters and their close association with strong air–sea fluxes means that these waters are most likely to be associated with heat content anomalies. Mode water anomalies have usually been studied at single locations with high density observations—for example, the Bermuda Atlantic Time series site (BATS; Talley 1996), or with repeat XBT sections, but full spatial distributions and total volume anomalies are difficult to determine from data. Formation theories for decadal mode water variability exist (e.g., Dickson et al. 1996; Curry and McCartney 2001) associated with atmospheric forcing distributions such as the North Atlantic Oscillation (NAO), but there is not yet enough observational data to clearly quantify the processes involved. Only a few modeling studies of variability in mode water production have been performed. Gulev et al. (2003) discussed variability of North Atlantic Mode Water formation in relation to surface fluxes and meridional heat transports from an ocean-only model. They showed that Labrador Sea and Subtropical Mode Water formation have the biggest response to the NAO, and found a 3-yr lagged response in the meridional heat transport. They did not look at the life cycle of the mode water anomalies within the basin or show that they were involved in the baroclinic adjustment leading to the 3-yr lag in heat transport. Marsh and New (1996) used some idealized forcing experiments in a North Atlantic isopycnal model to show variations in 18°C mode water formation in response to cold winters, and despite quantifying the anomalous volumes, they did not relate these to the forcing, or look at the evolution of these anomalies. In this paper we follow a more general method for analyzing water formation and transformation anomalies applicable to all water classes. Walin (1982) studied the relationship between sea surface heat flux and the thermal circulation of the ocean by analyzing the changes in volumes of water identified by their temperature. To address the relationship among mode water formation, subduction, and persistent heat anomalies, and following the work of Marshall et al. (1999), this diagnostic has been modified to include a measure of the subducted volume anomalies. The modified di-
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FIG. 1. The North Atlantic and Arctic Ocean region of the HadCM3 coupled climate model used in this paper for the analysis of water mass variability.
agnostic has been applied to 100 yr of control run data generated by the Third Hadley Centre Coupled Ocean– Atmosphere General Circulation Model (HadCM3), a fully coupled, free-running model without flux adjustment, which includes mixed layer physics that should reproduce realistic thermocline structures and ventilation processes. The only previous work on water mass diagnostics in coupled models (Speer et al. 2000) in a southern ocean study suggests that water transformations may be better represented in coupled models because of the more balanced mixed layers. This study looks at this issue for water masses in the North Atlantic and Arctic sector (see Fig. 1) of the HadCM3 model. The paper begins with a description of the model, including the mixed layer and water mass modes (section 2), followed in section 3 by a description of the Walin volume census diagnostic and how it is applied to the HadCM3 fields. The results of the analysis are split into two sections. In section 4, the 100-yr means are presented and discussed. These highlight the role of subduction in the formation of the main North Atlantic mode waters. Then the variability about these means is presented in section 5. Thermodynamic transformation pathways in the model are identified with lag-correlation analyses and discussed in the light of current theories and observations of ocean water masses. The paper is closed with a discussion in section 6.
2. The coupled HadCM3 climate model in the North Atlantic a. Model description HadCM3 is a global coupled atmosphere–ocean–sea ice climate model. The model runs freely without flux
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correction, therefore variability in the output fields represents the natural variability of the model system. The atmosphere consists of a 3.75° ⫻ 2.5° horizontal grid with 19 levels in the vertical. A full description of the atmospheric model can be found in Pope et al. (2000). The ocean consists of a 1.25° ⫻ 1.25° Arakawa B grid with 20 vertical levels, based on the primitive equation code of Cox (1984). There are parameterizations of isopycnal mixing and Gent and McWilliams (1990) adiabatic dissipation of potential energy. For every atmosphere grid cell, there are six ocean grid cells, and the land–sea masks of the model components match exactly to facilitate coupling. The sea ice model is based on Semtner’s (1976) zero-layer thermodynamic model and is run at the same resolution as the ocean model. Details of the ice model can be found in Gordon et al. (2000). The ocean component includes a hybrid scheme to parameterize the near-surface vertical mixing in which a Kraus and Turner (1967) mixed layer submodel is combined with a K-theory scheme (see Gordon et al. 2000). This produces realistic spatial and temporal variability in the mixed layer depth in the key formation regions. The background tracer diffusion coefficient is depth dependent ranging from 1 ⫻ 10⫺5 m2 s⫺1 at the surface to 5 ⫻ 10⫺5 m2 s⫺1 at 1500 m, with larger values below. This fits closely the theoretical/observational estimates of Kraus (1977). The ocean currents were initialized to zero, and initial temperature and salinity fields were taken from the Levitus and Boyer (1994) climatology. The atmosphere was initialized using an atmosphere-only integration. The ocean is integrated with a 1-h time step and the atmosphere with a 30-min time step. The ocean and atmosphere model components are coupled once per day.
b. Model control run and hydrography The model has been run for more than 1000 yr without significant drift in the model’s climate (global mean ocean temperature decreases by 0.03°C per century). This control run uses constant levels of trace gases based on preindustrial values, and incoming solar radiation provides the only external forcing. As with most models of this type, the Gulf Stream separates too far north; however, the ocean general circulation features are consistent with observations. The maximum strength of the North Atlantic meridional overturning circulation occurs at approximately 45°N and has an average value of 19 Sv (1 Sv ⬅ 106 m3 s⫺1; Cooper and Gordon 2002), which compares well with observations (e.g., Hall and Bryden 1982). The poleward heat transport across 24°N in the Atlantic is 1.0 PW (Cooper and
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Gordon 2002) compared with observational estimates of 1.2 PW (Hall and Bryden 1982) and 1.1 PW (Macdonald and Wunsch 1995). The breakdown of the poleward heat transport suggests that the model may be transporting more heat in the western boundary current and less heat through Ekman transport than estimated from observations (Gordon et al. 2000). This run also possesses low-frequency natural variability due to both ENSO (Collins et al. 2001) and NAO events (Collins et al. 2001; Cooper and Gordon 2002). The 100 yr of monthly mean fields made available for this study start from 590 yr into the integration. This period was chosen for its strong decadal NAO variation (Fig. 2a). There is a long-term cooling trend in the heat content of the North Atlantic Ocean, and the variation in the heat content about this trend (Fig. 2b) has a similar pattern to the NAO signal. Using the monthly mean fields it is possible to close the heat budget for the North Atlantic to within ⫾0.01 PW. This residual imbalance probably results from the missing highfrequency ⬘T ⬘ terms at the open boundary. These tend to be small at the equator; however, for smaller regions, the imbalance becomes more significant. Given the reasonable hydrography of the model, the good physical representation of the mixed layer processes, and the negligible drift without flux correction, this model should produce reasonably realistic water masses. The waters of the North Atlantic/Arctic Oceans were integrated into temperature–salinity (T– S) bins then converted to a percentage of the total volume (Fig. 3), where the bin widths are ⌬T ⫽ 0.5°C and ⌬S ⫽ 0.1 psu. The model is producing three distinct modes for waters above 11°C—these are a weak mode at 18.5°C, 36.5 psu; a stronger mode at 16.5°C, 36.3 psu; and a further strong mode at 14.0°C, 36.1 psu. Below 11°C there is a mode centered at 9.5°C, 35.7 psu, and another mode centered at 6.5°C, 35.7 psu. Two colder modes also exist. Bear in mind that the poor vertical resolution of the deeper model layers will restrict the variability of the deeper water mass properties considerably. The three warmer modes are formed on the equatorward boundary of the Gulf Stream extension/North Atlantic Current. The observed mode waters are the Subtropical Mode Water at 18.0°C, 36.5 psu (Worthington 1959; Talley and Raymer 1982); the Maderia mode waters formed between 16° and 18°C, 36.5 and 36.8 psu (Käse et al. 1989; Siedler et al. 1987); and the warmer range of the Subpolar Mode Waters at 10°–15°C, 35.5– 36.2 psu (McCartney 1982; McCartney and Talley 1982). Below 11°C the first mode represents the colder range of the Subtropical Mode Waters formed on the
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FIG. 2. (a) The North Atlantic Oscillation index for the 100 yr of the HadCM3 control run starting from year 590. The NAO index has been calculated using the definition of Hurrell (1995). The thin line shows the DJF index, and the thick line shows the data filtered as per Hurrell (1995). (b) The heat content anomaly for the North Atlantic and Arctic Oceans. The thin line shows the monthly anomaly and the thick line shows the data filtered using a low-pass Lanczos filter with a 5-yr cutoff.
poleward side of the North Atlantic Current; the second corresponds to the Labrador Sea Waters (LSW). Typically Labrador Sea Waters are observed to be the waters at 3.3°C, 34.8 psu (Talley and McCartney 1982); however, Cooper and Gordon (2002) have noted that
FIG. 3. T–S census of volumes in the North Atlantic–Arctic Oceans of HadCM3. The data represent the integrated volumes over the 100-yr period rather than the volumes of the mean fields. The T–S bins used are ⌬T ⫽ 0.5°C and ⌬S ⫽ 0.1 psu. A log scale is used for % water mass volume where the total volume in the basin is 1.8 ⫻ 1017 m3, and ⫹ points mark the main mode centers.
Fig 3 and 4 live 4/C
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the Labrador Sea waters formed in HadCM3 are too warm and too saline, and are strongly affected by mixing before the waters leave the Labrador sea. A feature of the T–S plot that stands out is the density banding in the colder water classes. This is associated with the discretization of volumes in the deep cells of the model, which have a vertical resolution of 615 m below 1500 m. It is apparent that the low vertical resolution in the deep waters is limiting the models ability to resolve the colder classes. Further consequences of this will be discussed later in the analysis. It was noted in the introduction that a good representation of the mixed layer depth variability is essential for producing realistic subduction rates. Subduction plays a key role in the analysis to be presented. The representation of the winter mixed layer base that will be used in the following analysis is defined as the local 100-yr maximum depth of the mixed layer, shown in Fig. 4. The key features are the trough beneath the Gulf Stream extension and the poleward side of the subtropical gyre (which is a little far north due to the Gulf Stream failing to separate correctly), the deepening of the mixed layer beneath the subpolar gyre, and the shoaling of the mixed layer into the subtropical region. It is this narrow band of mixed layer shoaling spanning the North Atlantic 25°–35°N that plays a key role in the ventilation of the subtropical thermocline. We now turn to the main water mass diagnostics that will be used for this study.
FIG. 4. The 100-yr maximum local winter mixed layer depth. This is the surface used to define the fixed internal surface WMLB in the diagnostics. Units are in m.
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3. Volume census diagnostic Walin (1982) presented a diagnostic for investigating the strength of the thermohaline circulation (THC) in terms of the thermodynamic transformation of water across potential temperature surfaces. Although later authors have preferred potential density coordinates, in this work potential temperature will be used to focus on the heat budget of the model. Thus wherever we refer to transformation and formation of water, we mean thermal transformation or formation, and this should be remembered when comparing with other work in the literature. The water volume is mapped onto a potential temperature coordinate, defining water volumes (masses) by their potential temperature. Following Walin’s (1982) analysis, the irreversible transformation of water across an isothermal surface toward colder temperature, within a bounded region, is given by G共T兲 ⫽
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dV共T兲 ⫺ 共T兲 共m3 s⫺1兲, dt
共1兲
where V(T ) is the volume of water with temperature less than T, and (T ) is the advective convergence at the open boundary of water with temperature less than T. Therefore G(T ) is, by convention, the transformation flux toward lower temperature. The two terms on the right-hand side of Eq. (1) can be calculated from the temperature field and the velocity normal to the open boundary. The transformation G(T ) can be further separated into surface transformations GS(T ) due to the action of surface fluxes at the outcropping of the isothermal surface, and internal transformations GI(T ) due to mixing and diffusion across the isothermal surface; that is, G共T兲 ⫽ GS共T兲 ⫹ GI共T兲 共m3 s⫺1兲.
共2兲
From a full knowledge of the surface fluxes and the sea surface temperature, it is possible to calculate GS(T ) using the equation
再
⭸ 1 ⫺ Q共T兲 GS共T兲 ⫽ ⭸T c
冎
⫺1
共m s 3
兲,
共3兲
where Q(T ) is the total heat flux in Watts into the sea surface where the sea surface temperature is less than T, and c ⫽ CP (J m⫺3 °C⫺1) is the heat capacity per unit volume. The internal transformation GI(T ) is calculated as a residual from the above equations. If the heat budget for the region is closed and there is no mixing or diffusion across the open boundary, then the residual is the transformation due purely to internal mixing. A simple test of heat budget closure is to integrate GI(T ) over all temperatures. This integral must be zero (no
mixing across the open boundary) as mixing and diffusion can only redistribute volume among the temperature classes, or equivalently, they cannot change the total heat content. The heat budget derived in this way for the diagnostics presented does close well for the North Atlantic as a whole. Two further diagnostics can be calculated from the transformations: 1) the formation rate, which is defined as the convergence of the transformation flux ⫺G/T. This diagnostic is used to show where water masses are being formed and removed in each thermal class. Then there is 2) the diathermal volume flux, which is calculated by integrating G(T ) with respect to temperature, and can be interpreted as the heat budget for the region of ocean concerned. This diagnostic is used to check the closure of the heat budget. When calculating these diagnostic components from the model fields, discrete temperature classes were used with a width ⌬T ⫽ 0.5°C. Tests show the main results are robust to this choice. The transformation volumes V(T ) are calculated using V共T兲 ⫽
兺V
ijk␦ijk;
␦ijk ⫽
i,j,k
再
1,
Tijk ⱕ T
0,
otherwise
.
共4兲
The volume convergence across the open boundary is given by
共T兲 ⫽
兺
GM 兲Aik␦ik; 共ik ⫹ ik
␦ik ⫽
i,k
再
1,
Tˆik ⱕ T
0,
otherwise
,
共5兲 where the caret indicates the spatially averaged temperature at the cell face and Aik is the area of the corresponding face. The ik are the normal components of velocity on the boundary, in this case at the equator. The integrated surface flux in Eq. (3) is given by Q共T兲 ⫽ ⫺
兺 S ␦ ; ij ij ij
ij
␦ij ⫽
再
1,
Tij ⱕ T
0,
otherwise
,
共6兲
where ij is the mean heat flux (in W m⫺2) into the horizontal grid cell (i, j) of surface area Sij. Provided a large number of cells are integrated when calculating the transformations, the functions will tend to become continuous. The volume time derivative in Eq. (1) is calculated as a finite difference between consecutive pairs of monthly averaged data. The boundary fluxes are timeaveraged between consecutive pairs of months so that they are aligned temporally with the volume time derivatives. Derivatives with respect to temperature are calculated as centered differences so that the temperature classes remain consistent with those used to calculate the transformation G(T ). It should be pointed out
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advection across the equator dominates the open boundary convergence.
a. North Atlantic water mass exchange and transformation
FIG. 5. (a) Mean diathermal volume flux and (b) mean water formation rate for the North Atlantic and Arctic Oceans calculated from 100 yr of HadCM3 control run monthly mean fields. The thick solid line shows trends over the 100 yr, x line is the surface heat flux contribution, dashed line is advection across the equator, and the o line is the mixing, which is calculated as a residual. The units are given on the axes.
that calculating the formation rate as a finite difference gives the mean formation rate between pairs of isotherms (T, T ⫹ ⌬T ), therefore these rates will vary with choice of ⌬T.
4. Mean transformations The diagnostics were applied to each set of monthly fields for the North Atlantic/Arctic Ocean region of HadCM3 (Fig. 1). From the time series of monthly diagnostics the 100-yr means were calculated. The mean transformations and formations are shown in Figs. 5a and 5b, respectively. The diathermal heat flux data have not been included, but they showed that the diagnostic closes the heat budget to the same order as noted in section 2. To make the figures easier to interpret, the temperature axis has been reversed so that a positive transformation moves volume to the right, as is conventional in potential density diagnostics (e.g., Speer and Tziperman 1992). Since formation is the gradient of the transformation, slopes in the transformation curves correspond to formation or destruction classes, and regions of zero slope correspond to classes where there is no build up or removal of waters due to the considered process. There are two open boundaries to this region, the equator and Bering Strait. In the control run the barotropic flow through Bering Strait is set to zero leaving only a small baroclinic component of flow so that
Under steady-state conditions the advective exchange across the equator also quantifies the net effect of transformations taking place within the basin. In the North Atlantic, waters are transformed from a warm surface inflow into a deep cold outflow that maintains the thermohaline overturning circulation. Schmitz and McCartney (1993) estimated the Atlantic equatorial volume exchanges within temperature classes (given in Table 1) using a variety of different datasets. They show an inflow of warm surface and thermocline waters (T ⬎ 7°C) at a rate of 13 Sv, and an inflow of cold (T ⬍ 1.8°C) Antarctic Bottom Waters (AABW) of 4 Sv, and therefore an outflow of about 17 Sv of North Atlantic Deep Waters (NADW) in the range 1.8°C ⬍ T ⬍ 7°C (neglecting Bering Strait and net precipitation contributions). The advection curve in Fig. 5a shows the equivalent THC equatorial transports calculated for HadCM3. Slightly different temperature ranges reflect the temperature class boundaries between inflow and outflow at the equator in HadCM3. Between 7.5° and 28°C there is a 15-Sv inflow of surface and thermocline waters, and between 1° and 2°C a 9-Sv inflow corresponding to the AABW. These two inflows are balanced by a net outflow of 24 Sv between 2° and 7.5°C corresponding to the NADW and intermediate waters (see Table 1). The upper-thermohaline transports in the model thus agree well with Schmitz and McCartney (1993) but the deep circulation appears to be too strong, by about 5 Sv. Surface heating and mixing work together to convert the inflow of warm surface waters to an outflow of cold NADW. The transformations due to surface heat fluxes (Fig. 5a) are warming waters with T ⬎ 25°C and cooling waters with T ⬍ 25°C. In contrast the mixing is cooling waters with T ⬎ 25°C and warming waters with T ⬍ 5°C, but at intermediate classes the mixing varies between warming and cooling. For the warmest waters (T ⬎ 25°C) there is a peak surface transformation of 22 Sv across the 28°C isotherm (Fig. 5a). The model is a little too warm in the Gulf of Mexico, which accounts for some of the highest temperatures. Surface processes are warming the waters and are directly balanced by mixing which transforms the waters back to colder classes at the same rate, also illustrated by the formation diagnostic (Fig. 5b). These temperature classes correspond to waters in the Tropics and subtropics where the isothermal surfaces
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TABLE 1. Comparison between observed and modeled North Atlantic equatorial transports associated with the thermohaline circulation. Observations*
HadCM3
Water mass
Temperature range (°C)
Transport (Sv)
Temperature range (°C)
Transport (Sv)
Surface layer Thermocline Intermediate NADW AABW
⬎24 7–24 4–7 1.8–4 ⬍1.8
⫹7 ⫹6 0 ⫺17 ⫹4
⬎25 7.5–25 5–7.5 2–5 ⬍2
⫹3 ⫹12 ⫺6 ⫺18 ⫹9
* Observation data and water mass definitions taken from Schmitz and McCartney (1993).
are essentially horizontal and therefore the diathermal mixing mainly moves water vertically. This balance between the surface and mixing transformations can be interpreted as the surface heating in the Tropics driving coldwater upwelling at a rate of 22 Sv. Table 1 suggests that perhaps there is too little equatorial inflow of the warmest waters in HadCM3 and therefore the heating transformations inside the basin in the Tropics may be on the large side to compensate. The intermediate temperature classes (Fig. 5a) show surface cooling peaks at 18°C (12 Sv) and 10.5°C (13 Sv), with an intervening minimum cooling at 14°C (6 Sv). Between 25° and 18°C, waters are increasingly cooled and destroyed by the surface transformations. The volume removed is in excess of the volume advected in across the Equator, with the balance provided by mixing, which forms waters in this temperature range by warming, that is, with heat supplied from the warmer waters (T ⬎ 25°C). Between 18° and 14°C the surface transformations lead to the thermodynamic formation of 6 Sv of water by cooling and there is also 3 Sv of inflow at the equator in this temperature range. We will refer to this range as the HadCM3 equivalent of the North Atlantic Subtropical Mode Water (STMW). Mixing removes STMW by warming waters with T ⬎ 16°C and cooling waters with T ⬍ 16°C. The STMW that are warmed replace waters lost through surface cooling between 25° and 18°C, suggesting an equatorward movement of STMW (see transformation pathways in section 5). The waters cooled by mixing become Subpolar Mode Waters (SPMW) in the range 10°C ⬍ T ⬍ 14°C. Between 10° and 7.5°C the surface transformations form about 1 Sv of water, and inflow at the equator provides 3 Sv, with mixing removing these waters by cooling to form colder classes. None of the classes down to 7.5°C contribute to the NADW outflow of the basin. Between 7° and 5°C the surface fluxes form about 3.5 Sv of water. These are the Labrador Sea Waters (LSW) of the model, which are about 2°C warmer and more saline than observations, as discussed in (Cooper and
Gordon 2002). Mixing produces another 3.5 Sv between 7° and 5°C and these become the warmest waters exported from the North Atlantic. Between 5° and ⫺1°C surface fluxes produce a constant cooling transformation—that is, they neither form nor destroy volume. However, mixing produces 8 Sv of water at 4.5°C, and 5 Sv at 3°C, which is exported from the basin. The total export at 3°C amounts to 9 Sv, with the additional 4 Sv of loss representing a consistent volume trend over the 100 yr, which shows up in Fig. 5b. Thus the total export of NADW from the North Atlantic amounts to 24 Sv, which appears in Table 1. The distinct mixing signatures defining the outflow NADW at least partly reflect processes on deeper model layers, which are poorly resolved, and yet have a dominant influence on the properties by the time the waters reach the Equator. Budgets for high latitude regions alone indicate a broader mixing formation across all these NADW classes. Between 2° and 1°C there is an inflow of 9 Sv of Antarctic Bottom Waters (AABW), 5 Sv of which are converted by mixing into 3°C water (see above); the remaining 4 Sv build up and appear as a trend in Fig. 5b. Thus the model is steadily cooling by gaining 2°C water while loosing 3°C water, which contributes to the global mean cooling trend of 0.03°C per century in the HadCM3 control run. The excess production and import of AABW may reflect inaccurate parameterisation of the convection and/or the sea ice processes in the Southern Ocean. For T ⬍ ⫺1°C the surface fluxes are forming polar waters in the Arctic basin at a rate of 9 Sv, which are then mixed back to warmer classes, particularly as the water exits the Arctic basin into the North Atlantic.
b. Mode waters and subduction To focus on the mode water classes the North Atlantic basin waters are separated about a surface representing the 100-yr local maximum winter mixed layer base (hereafter WMLB; Fig. 4). This surface divides the waters that may be in annual contact with the surface
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FIG. 6. Mean diagnostic curves calculated for the volumes of the North Atlantic and Arctic Oceans above and below the 100-yr maximum winter mixed layer base (WMLB). The volume convergence has been separated into the open boundary and WMLB components. (a) Mean diathermal volume fluxes above the WMLB, (b) mean water formation above the WMLB, (c) mean diathermal volume flux below the WMLB, and (d) mean water formation rate below WMLB. The thick solid line is the trend over 100 yr, the x line is the surface heat flux contribution, the dashed line is the advection across the equator, the o line is the internal mixing, and the ⫹ line is the subduction across the mixed layer base. The units are given on the axes.
from those that are definitely not (see Marshall et al. 1999). The subduction is defined as the convergence of water through the WMLB. Figures 6a–d show the mean transformation and formation processes for the HadCM3 North Atlantic, above and below the WMLB. Open boundary (equator) and WMLB convergence are shown separately as
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advection and subduction, respectively. Note that these two terms cancel at the highest T classes (Figs. 6a,c) as the net convergence must go to zero. Transformations above the WMLB in Fig. 6a show that subduction has broadly the same two peaks, around 19° and 10°C, as the surface transformation, with the subduction peaks being slightly larger. The extra mixed layer cooling is supplied by mixing across the WMLB at an average rate of 3 W m⫺2 between 20° and 5°C. Only where subduction does not play a role, for the warmest (coldest) water classes, does mixing lead to substantial net cooling (warming) of the mixed layer. Formation above the WMLB in Fig. 6b shows peaks associated with subduction that are much sharper than the surface formation peaks and focused around mode water classes. Subduction peaks are at 18°C (6 Sv), 7.5°C (7 Sv), and 2.5°C (9 Sv). The formation rates are based on the transformation changes 19°–17°C, 8°– 6.5°C, and 5°–1°C, respectively. The first two peaks correspond to the Subtropical Mode Waters and Labrador Sea Waters, respectively; the last is associated with water formed in the Nordic Seas. The Subpolar Mode Waters are in the range of a broad net obduction peak (negative subduction; Qiu and Huang 1995) around 12.5°C (5 Sv between 12° and 15°C) consistent with a net upwelling of water into the mixed layer in the subpolar gyre. For the 18°C waters, mixing transformations inside the mixed layer sharpen the properties of the waters prior to subduction, without significantly changing the total formation rate calculated using surface fluxes alone. Although mixing breaks the direct connection between surface formation and subduction at the same temperature class, this result gives some support to the Speer and Tziperman (1992) method of using surface transformations as an indication of the total production of 18°C water, at least if one allows for some shifting of the class range being examined. However the colder subduction peaks around 7.5° and 2°C are hardly identifiable at all in the surface transformations and are virtually entirely supported by mean formation of waters by mixing within the mixed layer. It would clearly be difficult to infer anything about subduction in these two classes directly from air– sea forcing, at least from the integrals over the whole basin. For the region below the WMLB Fig. 6d shows that mixing takes the waters from these two subduction peaks and forms intermediate (3°–6°C) water masses. It is the water masses within this temperature range that are exported from the North Atlantic basin as NADW. Figure 6a shows that mixing forms 7.5°C waters by
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cooling waters warmer than 7.5°C and warming waters colder than 7.5°C. Figure 6c shows that, once below the WMLB, water with T ⬎ 8°C is mixed up to the warmer SPMW where it is obducted back into the mixed layer, while water with T ⬍ 8°C is mixed down to form the NADW, which exits the basin at around 6°C. Figure 6b also shows that water is in a net sense entraining, or obducting, back into the mixed layer around 21° and 12°–14°C. Figures 6c and 6d suggest a link between the subducted 18°C and the obducted 21°C water via mixing below the WMLB. Between 16° and 22°C mixing is warming waters with the subducted 18°C STMW being mixed into 21°C waters that are obducted back into the mixed layer. This thermodynamic pathway for the recirculation of the STMW is examined further through time lagged correlations in section 5. Waters between 12° and 14°C correspond to the warmer range of the SPMW. As water is advected eastward and northeastward around the subpolar gyre it encounters progressively deepening winter mixed layer depths and is progressively cooled. Here the surface forcing is acting to destroy these water masses, which are resupplied by obduction, and some mixing, across the WMLB (Fig. 6c). The formation diagnostic below the WMLB (Fig. 6d) shows there are two sources of these obducted waters. Between 12° and 14°C the waters are advected in through the open boundary, and between 11° and 12°C the waters are formed by the warming of 9°C waters through mixing. This temperature range indicates where water mass anomalies below the mixed layer are able to return to the surface allowing feedback with the atmosphere.
5. Natural variability It has been shown above that the mean North Atlantic water transformations in HadCM3 are broadly consistent, both qualitatively and quantitatively, with interpretations of the hydrographic record, particularly in regard to the model mode waters. Two different methods of studying the variability in water mass transformations are presented. First anomalies about the mean transformations in Fig. 5a are presented as Hovmoeller diagrams on a temperature versus time plot. These provide information on the magnitude of the variability, the associated time scales, and whether any classes are more important than others. Time in the diagrams is defined in years from the start of the 100-yr dataset. The second method uses lag correlations between pairs of transformation processes to identify causal pathways. These tell us something of the processes driving
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formation and destruction of anomalous water masses, and the time taken for processing pathways to operate. In all the diagnostics the monthly transformations were filtered using a 5-yr Lanczos filter to remove the high frequency variability, particularly the dominant seasonal signal, in order to reveal longer-term changes that are most relevant to climate.
a. Total transformation variability The anomalous transformations about the means of the four components in Fig. 5a are presented in Figs. 7a–d, along with their corresponding 100-yr means below each figure. A negative anomalous transformation corresponds to increased transformation toward warmer temperatures, representing a positive heating anomaly. The variability in advection across the equator (Fig. 7c) is everywhere smaller than 2 Sv (with the exception of the AABW classes); therefore, the mismatch between total and surface forced transformations is mainly accounted for by mixing. Hence variations in water volumes in the North Atlantic thermocline are not generally exported. The fact that the variability due to advection is small implies that the anomalous volume transformations in Fig. 7a represent G⬘(T ) and are equal to the sum of the anomalous surface and mixing transformations. In the warmest classes (T ⬎ 20°C) there is very little total anomalous transformation (⬍2 Sv). However, both the surface and mixing anomalies (Figs. 7b,d) show large variability, with the mixing anomalies canceling the surface anomalies. Around 28°C the anomalies tend to form as dipoles representing a shift in temperature of the transformation peak in Fig. 5a. The correlation between the anomalous surface and mixing transformations, shown in Fig. 8, confirms that these two processes are strongly anticorrelated for all classes above 20°C. Mixing responds quickly to surface flux variability so that no anomalous water masses build up in these classes. This is a region where water is mainly returned from the thermocline into the mixed layer (see Fig. 6c) so waters that are formed by surface fluxes remain within the mixed layer. Between 10° and 20°C, corresponding to the STMW and SPMW ranges, Fig. 8 shows that the volume transformation anomalies are most highly correlated with the surface anomalies. The STMW classes vary on a 5-yr time scale with peak anomalies less than 5 Sv. The SPMW classes vary on 5–10-yr time scales, with peak anomalies around 6 Sv. Figures 7a and 7b clearly indicate surface forcing as the main cause of the total transformation variability in these classes. For waters below 10°C there are strong positive cor-
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FIG. 7. Hovmoeller plots in temperature and time coordinates of the transformation anomalies about the 100-yr means for the North Atlantic and Arctic Oceans of HadCM3. (a) Total transformation anomalies, (b) surface forcing transformations, (c) open boundary advection anomalies, and (d) mixing transformation anomalies. Each figure includes the associated mean curve. For each plot the data have been filtered using a low-pass Lanczos filter with a 5-yr cutoff. Units for the anomalies are Sverdrups.
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FIG. 8. Correlations between the filtered anomalous total transformations (V/t), anomalous surface transformations (GS), and anomalous mixing transformations (GI) for each potential temperature class of the North Atlantic/Arctic Oceans of HadCM3.
relations between the total transformations and mixing. This is consistent with Fig. 6b, which shows that for 0 ⬍ T ⬍ 10°C mixing is the dominant mechanism of mean water formation when defined in temperature classes. Below 0°C the surface and mixing transformations are again strongly anticorrelated in Fig. 8, showing that Arctic waters are formed directly by surface cooling. To investigate the development of anomalies and to follow their transformations the processes above and below the WMLB are focused on using lag correlations in the following section.
b. Transformation pathways—Subtropical Mode Waters Figure 9a shows the correlations between subduction and surface forced transformations in all classes, for lead and lag times up to 6 yr. The shaded regions identify where the correlations are statistically significant to a 95% level of confidence, calculated by assuming 20 independent samples based on the filter width of 5 yr. The two features that stand out are (i) surface forcing leads subduction in the 18°C STMW class by about 2 yr and (ii) subduction leads surface forcing, also by about 2 yr, for SPMW at around 12°C. In both cases peak correlations are above 0.6. The SPMW case will be discussed in the following subsection while the focus here is on the STMW. After being subducted the STMW anomalies can be followed by lag correlating the anomalous transformations at one temperature class (where STMW are formed) against transformation anomalies in all other temperature classes, for the three key processes of subduction, mixing and surface forcing. Figure 9b shows subduction at 18.5°C lag correlated against subduction at all other temperature classes. Subduction at 21°C is strongly negatively correlated (obduction) with a lag of about 3 yr. A similar pattern but of opposite sign is seen
FIG. 9. (a) Subduction across the winter mixed layer based lag correlated against the surface forced transformations in the mixed layer. The two thick vertical lines indicate the peak classes associated with the STMW (18.5°C) and the SPMW (12.5°C). (b) Subduction across the mixed layer base at 18.5°C lag correlated against subduction at all other temperatures. The thick vertical line highlights the 18.5°C temperature class. The contour interval is 0.2 with solid lines indicating positive and zero correlations and dashed lines negative. The data were low-pass filtered prior to the correlation calculation using a Lanczos filter with a 5-yr cutoff. The gray shaded regions show areas where the statistical significance is ⬎95%, calculated assuming that at most there are 20 independent samples due to the filtering.
for mixing below the WMLB (not shown) indicating a mixing dipole to higher temperatures with a 3-yr lag. Further investigation of mixing within the mixed layer shows that after obduction at 21°C the water anomalies are immediately dispersed (no lags) without detectable feedback on surface transformations. Thus a typical sequence of events might be (a) anomalous surface production of STMW, (b) subduction of this mode water 2 yr later, (c) mixing of this water with warmer water classes below the WMLB over the following 3 yr, and (d) anomalous obduction of warmer waters back across the WMLB 3 yr after the original subduction. The whole sequence lasts 5 yr. The 2- and 3-yr lags, above and below the WMLB, respectively, must imply storage of anomalous waters, during which time the waters may be transported around the subtropical gyre. The general southwards movement in the gyre would be consistent with the dominant transformation of the subducted STMW, involving mixing with warmer water classes, prior to eventual reentrainment into the mixed layer. It should be noted that these lag correlations are basin averages and longer and shorter lags in the processing of STMW
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anomalies can exist in different regions, and are the subject of ongoing research. The dominant process adding noise to these signals is probably the wind stress variability, which may affect subduction/obduction rates, advection around the gyre, and also mixing through modifying temperature gradients.
c. Transformation pathways—Subpolar Mode Waters Subpolar Mode Waters tend to reside within the deep winter mixed layer of the subpolar gyre where they are overcapped each summer by a thin seasonal mixed layer. As they propagate around the northern North Atlantic they are cooled through contact with the atmosphere (McCartney and Talley 1984). The SPMW therefore have a broad range of temperatures, and follow three distinct pathways in their conversions. The warmest SPMW can be found east of Newfoundland near the start of the North Atlantic Current, where mixed layer waters from the subtropical gyre cross into the subpolar gyre and are advected in a northeast direction. Some of these waters cross back into the subtropical gyre and return southward in the east Atlantic (Siedler et al. 1987; Paillet and Arhan 1996), while the remainder turn north. This northward current then splits, with some of the water entering the Nordic Seas where it is cooled and returns as Denmark Strait overflow water, and the remainder of the water circulates westwards again, with some entering the Labrador Sea. The waters formed in the Labrador Sea and the Nordic Seas combine to form the NADW. Figure 9a shows that for SPMW around 12°C, subduction anomalies lead surface transformation anomalies by approximately 2 yr with a positive correlation, while Fig. 6b shows that 12°C water is on average obducting back into the mixed layer. Therefore it is expected that subduction anomalies should lead surface transformations, and also that such anomalies in obduction could potentially be a precursor to surface flux variability. Figure 10a shows the lag correlations between the surface transformations at 12.5°C (the dominant class of SPMW; McCartney and Talley 1982) and the surface transformations for all other classes. It can be seen that anomalous surface transformations from classes as warm as 19°C are positively correlated with the surface transformations at 12.5°C, with the lead times increasing up to 5 yr with increasing temperature. This is consistent with near-surface water mass anomalies advecting northward from the subtropical gyre into the subpolar gyre while being cooled by surface fluxes. This is similar to the transformation route described by McCartney et al. (1996) and Sutton and Allen (1997). Therefore, we find two main sources for anomalous
FIG. 10. (a) Surface transformation anomalies at 12.5°C, lag correlated against surface transformations at all other temperatures. (b) Subduction at 6.5°C, lag correlated against subduction at all other temperatures. (c) Surface transformations at 12.5°C, lag correlated against subduction at all other temperatures. The contour interval is 0.2 with solid lines indicating positive and zero correlations and dashed lines negative. The data were low-pass filtered prior to the correlation calculation using a Lanczos filter with a 5-yr cutoff. The gray shaded regions show areas where the statistical significance is ⬎95%, calculated assuming that at most there are 20 independent samples due to the filtering.
SPMW: advection of surface anomalies in the gyre system, and the obduction of anomalies into the mixed layer from the thermocline below. Both Curry and McCartney (2001) and Cooper and Gordon (2002) have suggested that Labrador Sea Water anomalies circulating deep within the subpolar gyre may affect the upper gyre circulation, and therefore precursors linking the models Labrador Sea waters and the obduction in SPMW classes were sought. Figure 10b shows subduction anomalies at 6.5°C (within the LSW range) lag correlated against subduction at all other classes. This indicates that LSW subduction precedes obduction variability for all SPMW classes (9°– 14°C) by 3–5 yr, with correlations over 0.6. The sense of the correlation is such that increased subduction of LSW leads to reduced obduction (increased subduc-
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tion) of SPMW 3–5 yr later. No significant correlations in conversions between LSW and SPMW classes below the WMLB were found; therefore, it is inferred that the correlations in Fig. 10b are dynamic in origin. This is consistent with the mechanism suggested by Curry and McCartney (2001) and Cooper and Gordon (2002). Figure 10c shows lag correlations of the anomalous surface transformations at 12.5°C against subduction in all other classes. The SPMW subduction signal leading surface transformations at 12.5°C by 2 yr, already seen in Fig. 9a, is again visible. Another significant signal shows surface transformations at 12.5°C leading the subduction of waters from 0° to 2°C by about 6 yr, with a negative correlation of about 0.5 (slightly lower than the previous correlations discussed). These cold temperature classes are associated with the Nordic Seas and Denmark Strait flows, and the correlation suggests additional surface cooling of SPMW is associated with cooling of the mixed layer by subduction in these colder classes some 6 yr later. Again there is no clear surface transformation or mixing pathway between these waters, and we infer that the connection is dynamical— that is, cooling of SPMW causing circulation changes, which change the subduction of cold waters in the Nordic Seas. Further investigations of these connections are needed to fully understand the causal links.
the same temperature. The normal 2-yr lead of surface forcing seen in Fig. 9a is quite noticeable, except around year 30, when the very large water anomalies in Fig. 11a are being formed. At this time surface forcing appears to lead the subduction by about 3 yr. This suggests how such a large anomaly might be built up in a region that is normally dominated by obduction. It appears that the anomaly in surface forcing at this time was strong enough to alter the obduction conditions and to produce large water mass anomalies below the mixed layer as a direct result of the surface forcing.
6. Discussion The main results of this paper are as follows: • a quantification of the thermodynamic processes
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d. NAO connections During the 100-yr period analyzed, the HadCM3 data showed remarkably large decadal NAO variations. Figure 2 indicates that there was a large heat anomaly associated with this NAO variability, and it is of interest to ask what water masses were involved. Figure 11a shows the integral of the total transformation flux from the beginning of the 100-yr period—that is, the total anomalous volume transformed. A positive (negative) value at a given temperature and time indicates the existence of an anomalous volume of water with warmer (colder) temperatures. Between years 30 and 55, while the NAO was persistently positive, there was around 1015 m3 of additional water with temperatures between 10° and 15°C (i.e., the SPMW range) present at the expense of colder waters. Most of this large volume anomaly was formed quickly within a few years of year 30 and thereafter it persisted below the WMLB for the next 25 yr. Comparing Figs. 7a and 7b suggests that surface forcing was responsible for the formation of this water mass, however it was also shown in Fig. 5b that SPMW classes are normally destroyed in a net sense by surface forcing. How can these results be reconciled? Figure 11b shows filtered time series of the subduction at 12.5°C and the surface forced transformations at
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driving the thermohaline circulation in the North Atlantic of the coupled model HadCM3, using potential temperature transformations, and some comparison with observations; quantifying Subtropical and Subpolar Mode Water formation and their subduction/obduction across the winter mixed layer base; demonstration of the natural life cycle of Subtropical Mode Water anomalies over periods up to 5 yr, and the transformation pathways involved; important lag correlations between processes controlling Subpolar and Labrador Sea Mode Waters were also found, but the complete life cycle of water mass anomalies could not be followed; and identification of decadal variability in Subpolar and Subtropical Mode Waters associated with NAO variations.
Overall the HadCM3 model produces fairly realistic representations of most North Atlantic water masses, apart from the Labrador Sea Waters that are too warm, and a realistic exchange with the rest of the global ocean. There is a 4–5-Sv excess import of Antarctic bottom waters leading to a slow cooling of the basin, and the NADW export at the Equator is strongly concentrated at 2.5°, 4.5°, and 6°C, due to only a few deep model layers. Subtropical Mode Waters are formed and subducted below the winter mixed layer base at about the same mean rate (6 Sv), which contrasts with the ocean-only model results of Marshall et al. (1999), who find much less net subduction than surface formation in these water classes, with mixed layer mixing always strongly opposing the effects of surface formation. It is possible that the close balance between surface formation and subduction of STMW in HadCM3 is due to it being a coupled model without flux correction. Ocean-only models typically use specified surface fluxes and a re-
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FIG. 11. (a) Total volume anomaly, relative to the start of the 100-yr dataset, of water above a given potential temperature. This is defined as the negative of the time integral of the total transformation anomaly. These data have been filtered using a low-pass Lanczos filter with a 5-yr cutoff. Included to the right of the figure is a plot of the corresponding NAO index (red) and the standardized total heat content anomaly (blue) for the North Atlantic–Arctic Oceans of HadCM3. (b) Subduction and surface transformation anomalies at 12.5°C presented as a waterfall plot in time. The offset between the two curves is set to the dominant lag time. The sloping lines are 1 to 1, highlighting the predominant 2-yr lag between subduction and surface transformation in this class. At year 30 it is apparent that the surface leads subduction by 3 yr. This is identified (see text) as the source of the long-term volume buildup.
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laxation to SST; thus mixed layer temperatures are determined by different, and possibly inconsistent, processes from those controlling deeper model layers. This could require substantial additional mixed layer mixing, destroying much of the water being formed by surface processes prior to subduction. Speer et al. (2000) first noted the importance of this inconsistency in a coupled model study of the Southern Ocean. Turning to the life cycle of Subtropical Mode Water anomalies, the 2-yr lag found between variability in surface formation and the correlated variability in subduction must represent the average time taken for water mass anomalies to be mixed between the surface and the winter mixed layer base (when defined as the 100-yr maximum). These delays also mean that water can potentially be subducted in a different location from where it is formed in the mixed layer. This is consistent with the Lagrangian view of subduction taken by Nurser and Marshall (1991) and Woods and Barkmann (1986), who note that buoyancy must be added following the Lagrangian motion of a water column (rather than in an Eulerian sense) in order for subduction to occur. After subduction the Subtropical Mode Water anomalies can be followed for another 3 yr as mixing anomalies convert them to warmer classes and they are obducted back into the mixed layer 5 yr after formation. We again expect the 3-yr delay to be associated with advection of waters around the subtropical gyre. Such a circulation of water mass anomalies is consistent with the ventilated thermocline view of the subtropical gyre presented by Luyten et al. (1983); however, our results emphasize that, at least in HadCM3, there is significant mixing and transformation during the circulation around the gyre, prior to the water re-entraining into the surface mixed layer. Also the 3-yr period is only a basin mean, and longer and shorter lifecycles of mode water anomalies may occur. This efficient conversion of mode water anomalies demonstrates how the mean subduction of 6 Sv of Subtropical Mode Waters is transformed away within the HadCM3 thermocline. The Subpolar Mode Waters in the model (10°C ⬍ T ⬍ 14°C) are dominated by obduction of water into the surface mixed layer in the subpolar gyre. In the mean about 5 Sv are entrained into the mixed layer, peaking around 12.5°C, with a further 2 Sv being formed by mixing around 11°–12°C, and these waters are then destroyed by air–sea forcing, primarily by cooling as water is advected northward and eastward around the subpolar gyre. Anomaly correlations reveal influences from the subtropical gyre, with leads of up to 5 yr, as warmer surface anomalies are advected across into the subpolar gyre, producing accompanying surface transformation
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anomalies. In addition, variations in obduction from below the mixed layer base lead surface transformations by about 2 yr, reflecting mixing between the base of the winter mixed layer and the surface. The existence of these precursors to surface transformation suggests that the subpolar gyre could be an area where ocean water mass anomalies may be precursors to atmospheric variability, although this has not been demonstrated. Also the upwelling and return of waters to the surface from below is consistent with the Samelson and Vallis (1997) model of the subpolar thermocline as being dominated by diffusive/mixing processes. Certainly water masses with ⫺1°C ⬍ T ⬍ 10°C are dominated by mixing formation. For example, surface formation of Labrador Seawater (5°C ⬍ T ⬍ 7°C) occurs at 3.5 Sv, but the mean subduction is 7 Sv, at slightly warmer classes (6.5°C ⬍ T ⬍ 8°C), with the difference made up by mixing formation. Nordic Sea Water has a subduction of 9 Sv from (1°C ⬍ T ⬍ 5°C) due almost entirely to mixing above the winter mixed layer base. The surface formation (9 Sv) at ⫺1.8°C, presumably associated with ice processes, is the necessary water source for the production of all of these other cold water masses, via mixing. When variability was examined, Labrador Sea Water subduction anomalies at 6.5°C lead obduction variability in SPMW by 3–5 yr, supporting the suggestion by Curry and McCartney (2001) of a relationship between Labrador Sea waters and the upper-ocean response in the region of the North Atlantic Current region. This connection appears to be dynamic, as we found no evidence for direct conversions between Labrador Seawater anomalies and SPMW classes below the mixed layer base. There is also some evidence that surface transformations of SPMW anomalies lead the subduction of Nordic Sea waters by about 6 yr, although the correlations involved are lower, and hence less significance. Finally the large ocean heat content anomaly associated with the unusual multidecadal persistence of positive NAO conditions was shown to consist mainly of SPMW in the subpolar gyre, replacing colder waters below. These SPMW were formed by surface formation within a few years during the transition to positive NAO conditions, reversing the normal pattern of surface destruction of this water mass class. Further investigations are really needed to understand more fully the origin of such extreme water mass events in the coupled system. The limitations of this study come from several areas. The HadCM3 model is fairly low resolution, particularly in the vertical, and this places a strong limitation on the preservation of water masses particularly in deeper layers. Numerical mixing is also likely to be
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significantly influencing water mass properties over time. The lack of mesoscale eddies due to the low horizontal resolution is less likely to qualitatively affect the results, although eddies do have a significant role affecting stratification near the mixed layer base (Marshall 1997), so they could influence the relationships and time scales associated with subduction found in this paper. The use of potential temperature diagnostics to define water masses has the advantage that it is conserved (unlike potential density) in the presence of mixing. The disadvantages are that colder water classes are not strongly separated by stratification and therefore the effects of internal mixing are likely to be dominant (e.g., in Fig. 8 mixing drives most of the total formation below about 10°C). In addition the diathermal volume flux is not a virtual mass flux as in buoyancy and potential density diagnostics. Care must also be taken when comparing transformation rates in potential temperature, with published results using potential density, because the variable thermal expansion coefficient makes a big difference to the results. For the future the spatial structure of the water mass anomalies must be explored more fully. The lag correlations found for the basin as a whole suggest that there would be time for significant movements of water as the transformation pathways are proceeding. Preliminary results indicate that the spatial development of the water mass anomalies can be followed, and these results will be presented in a future paper. It will also be interesting to seek atmospheric responses that lag predictable water mass variations, leading to potential climate predictability, realizable through observation and assimilation of anomalous water masses into coupled models, possibly using ensemble methods. Acknowledgments. This work was supported by NERC under the COAPEC thematic program. Thanks go to Bablu Sinha for his assistance in determining the model fields required to close the heat budget and for his useful discussions of this work. Many useful discussions were also had with George Nurser, Jonathan Gregory, and members of the COAPEC team. REFERENCES Collins, M., S. F. B. Tett, and C. Cooper, 2001: The internal climate variability of HadCM3, a version of the Hadley Center coupled model without flux adjustment. Climate Dyn., 17, 61–81. Cooper, C., and C. Gordon, 2002: North Atlantic oceanic decadal variability in the Hadley Centre coupled model. J. Climate, 15, 45–72. Cox, M. D., 1984: A primitive equation, 3 dimensional model of the ocean. GFDL Ocean Group Tech. Rep. 1, 143 pp.
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