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JOURNAL OF CLIMATE—SPECIAL SECTION
VOLUME 19
North Atlantic Subtropical Mode Waters and Ocean Memory in HadCM3
CHRIS OLD
Institute of Atmospheric and Environmental Science, University of Edinburgh, Edinburgh, United Kingdom
KEITH HAINES
Environmental Systems Science Centre, University of Reading, Reading, United Kingdom
(Manuscript received 23 January 2005, in final form 15 August 2005)
ABSTRACT
A study of the formation and propagation of volume anomalies in North Atlantic Mode Waters is
presented, based on 100 yr of monthly mean fields taken from the control run of the Third Hadley Centre
Coupled Ocean–Atmosphere GCM (HadCM3). Analysis of the temporal and spatial variability in the
thickness between pairs of isothermal surfaces bounding the central temperature of the three main North
Atlantic subtropical mode waters shows that large-scale variability in formation occurs over time scales
ranging from 5 to 20 yr. The largest formation anomalies are associated with a southward shift in the mixed
layer isothermal distribution, possibly due to changes in the gyre dynamics and/or changes in the overlying
wind field and air–sea heat fluxes. The persistence of these anomalies is shown to result from their subduction beneath the winter mixed layer base where they recirculate around the subtropical gyre in the
background geostrophic flow. Anomalies in the warmest mode (18°C) formed on the western side of the
basin persist for up to 5 yr. They are removed by mixing transformation to warmer classes and are returned
to the seasonal mixed layer near the Gulf Stream where the stored heat may be released to the atmosphere.
Anomalies in the cooler modes (16° and 14°C) formed on the eastern side of the basin persist for up to 10
yr. There is no clear evidence of significant transformation of these cooler mode anomalies to adjacent
classes. It has been proposed that the eastern anomalies are removed through a tropical–subtropical water
mass exchange mechanism beneath the trade wind belt (south of 20°N). The analysis shows that anomalous
mode water formation plays a key role in the long-term storage of heat in the model, and that the release
of heat associated with these anomalies suggests a predictable climate feedback mechanism.
1. Introduction
The storage and transport of heat by the ocean contributes significantly to decadal climate variability
(Rossby 1959; Hasselmann 1991; Deser and Blackmon
1993; Rodwell et al. 1999; Seager et al. 2000; Marshall et
al. 2001; Wu and Gordon 2002; Levitus et al. 2005). In
turn, variability in the atmospheric processes drives
long-term changes in ocean heat content and ocean
heat transport (Molinari et al. 1997; Curry et al. 1998;
Seager et al. 2000). The key difference between the
ocean and the atmosphere is the time scales over which
each responds to change. The atmosphere produces fast
Corresponding author address: Dr. Chris Old, Institute of Atmospheric and Environmental Science, University of Edinburgh,
The Crew Building, The King’s Buildings, West Mains Road, Edinburgh EH9 3JN, United Kingdom.
E-mail: [email protected]
© 2006 American Meteorological Society
process responses, while the ocean, because of its
greater inertia, viscosity, and heat capacity, responds
slowly to system perturbations.
Observations and modeling studies show that temperature anomalies formed off the east coast of North
America propagate along the North Atlantic Current
(Hansen and Bezdek 1996; Sutton and Allen 1997).
Krahmann et al. (2001) attributed the formation of
these anomalies to the dominant mode of atmospheric
variability in the North Atlantic, the North Atlantic
Oscillation (NAO). Eshel (2003) suggests that this thermal persistence and coherent evolution is consistent
with the subtropical gyre playing an active role in North
Atlantic climate variability. Sturges et al. (1998) showed,
through a modeling study, how decadal wind forcing
could produce the observed variability in the subtropical gyre. Inui and Liu (2002) argued, using the results
from a process model, that midlatitude wind-forcing
effects can form subducted temperature anomalies.
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OLD AND HAINES
An alternative perspective to considering temperature anomalies is to consider the thickness and volumes
of the water masses themselves. Decadal variability observed in the 18°C subtropical mode waters (STMWs)
of the North Atlantic has been attributed to the NAO
(Joyce et al. 2000). McCartney (1982) made the point
that subtropical mode waters are not simply locally
formed water masses, but that lateral advection carries
them away from their formation region. Their persistence results from their low potential vorticity (i.e.,
large thickness) at formation, and subsequent equatorward transport into regions where they lie beneath the
seasonal mixed layer base. Talley and Raymer (1982)
identified 18°C mode water anomalies propagating
around the subtropical gyre on time scales of 5 yr, with
variability in the 18°C water properties over 5 to 10 yr.
The observed thickness anomalies were not directly related to changes in the surface fluxes in the formation
region; it was suggested that either changes in the wind
stress or the advection of anomalies from outside the
formation region were the most probable sources of
variability.
Schneider et al. (1999) observed the propagation of a
temperature (volume) anomaly in a long-term record of
upper-ocean hydrography of the North Pacific. They
showed the formation and subduction of an anomaly in
the eastern North Pacific and its subsequent propagation in the mean geostrophic flow around the subtropical gyre. The anomaly took 8 yr to propagate from the
formation region down to the tropical Pacific. Zhang et
al. (1998) showed a strong connection between the
propagation of such heat anomalies into the Tropics
and the decadal variability in El Niño.
The implication is that mode waters, through their
preferential formation, low potential vorticity, and
rapid subduction into the permanent thermocline, can
store heat anomalies that result from atmospheric variability on interannual to interdecadal time scales. The
spatial and temporal distribution of the long-term hydrographic record is sufficient for analyzing bulk
change (e.g., Levitus et al. 2005). However, deficiencies
in the record make it difficult to resolve the subbasinscale details of mode water evolution in response to
climatic variability from the observations alone. More
recently, high-resolution GCMs have been used to analyze mode waters (New et al. 1995; Marshall et al. 1999;
New et al. 2001; Gulev et al. 2003). These ocean-only
models were driven by seasonal climatologies of sea
surface heat flux and wind stress, and the work aimed to
better define mean mode water formation processes,
subduction, and renewal rates. To date, very little work
has focused on determining the time scales associated
with variability in formation and the persistence of
1127
anomalous mode waters. Quantifying these two time
scales will help determine whether or not mode waters
play a key role in the long-term evolution of oceanic
heat anomalies, and their significance in climate
change.
Both Speer et al. (2000) and Haines and Old (2005)
suggest that water formation and subduction in coupled
models (and by implication in the real ocean) may vary
differently to that in forced ocean-only models because
of the complexities of atmosphere–ocean feedbacks.
Haines and Old (2005) noted that, in the coupled
HadCM3 model, the role of the mixed layer in buffering subduction from surface water formation was reduced for the subtropical mode waters. This allows
more of the variability in surface water formation to be
subducted, leading to persistent mode water anomalies
below the winter mixed layer base. This increase in
subduction variability is likely to be critical for understanding the life cycle of mode water anomalies; hence
this study is most naturally carried out in the framework of a coupled climate model.
The work presented in this paper forms part of an
ongoing study of the evolution of North Atlantic oceanic heat anomalies in the Third Hadley Centre
Coupled Ocean–Atmosphere GCM (HadCM3), developed by the Met Office. HadCM3 has been run freely
for more than 1000 yr without flux correction. This implies that the variability observed in the model data is
natural to the system. The model also includes mixed
layer physics, which plays a key role in the correct formation of mode waters (see below). The analysis presented is based on 100 yr of monthly mean fields taken
from the 1000-yr control run dataset. A description of
the model and data used is given in the paper by Haines
and Old (2005), which presents a basin-integrated diagnostic study of the water mass variability in the North
Atlantic region of HadCM3.
Haines and Old (2005) showed that, in the 100-yr
dataset, there is significant variability in the North Atlantic water masses on pentadal to decadal time scales,
and that this variability can be attributed to a subset of
water masses identified as mode waters. In particular it
was shown that, in an integrated sense, volume anomalies of 18°C water subducted into the permanent thermocline are strongly lag correlated with mixing to
warmer classes, followed 3 yr later by obduction from
the thermocline back into the mixed layer. The interpretation given was that volume anomalies in the 18°C
mode water propagate around the subtropical gyre, mix
to warmer classes then obduct back into the mixed
layer 5 yr (on average) after their initial formation.
However, the integrated diagnostic gives no spatial information, and therefore this interpretation needs to be
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JOURNAL OF CLIMATE—SPECIAL SECTION
verified through a study of the spatial variability of the
associated volume anomalies.
The focus of this paper is the mode waters in the
subtropical North Atlantic. To emphasize the relationship between volume anomalies and ocean heat content, the analysis is based on water masses defined by
their potential temperature. This will be shown to be
reasonable for the warmer classes of water masses considered here. An analysis based on thickness anomalies
between pairs of isotherms is used to detect propagating volume anomalies. The key questions to be addressed in this paper are as follows: What are the time
scales of variability associated with mode water formation? How long do mode water anomalies persist and
how does their spatial distribution relate to variability
in formation? What is the role of mode water in oceanic
heat content variability? The paper begins in section 2
with a brief description of subtropical mode waters and
their formation, where the relevant fields from the
model are presented and discussed. In section 3 the
formation and subduction regions for the subtropical
mode waters are identified and analyzed using spatial
maps of the thickness variance. In section 4 the persistence of mode water anomalies is determined using spatial maps of thickness anomalies lag correlated against
the variability in the subduction region. The results of
the analyses are discussed in section 5, with emphasis
on the role of mode waters in anomalous ocean heat
storage.
2. North Atlantic Mode Waters in HadCM3
a. Warm North Atlantic Mode Waters
Hanawa and Talley (2001) identify five different generic types of mode waters that are observed around
the world; of these, the first three are found in the
subtropical North Atlantic. The main type of STMW
found in all ocean basins is that associated with western
boundary current extensions. In the North Atlantic this
corresponds to Worthington’s (1959) 18°C waters,
which have an average salinity of 36.5 psu and a potential density of 26.5. The formation region for this mode
is just south of the Gulf Stream extension, at approximately 65°W, in an area of high surface heat loss to the
atmosphere. This water spreads southward into the
subtropics beneath the subtropical gyre, extending to
45°W and as far south as 20°N (McCartney 1982).
The second type of mode water is that formed in the
eastern part of the subtropical gyres. In the North Atlantic this corresponds to the Madeira Mode Water
(MMW). The equatorward side of the Azores front and
the western side of the Canary Current bound this
mode water. Käse et al. (1985) and Siedler et al. (1987)
VOLUME 19
define the Madeira mode water as having a potential
temperature between 16° and 18°C, salinity between
36.5 and 36.8 psu, and a potential density between 26.5
and 26.8. This mode water was found by Siedler et al.
(1987) to be relatively short lived with anomalies tending to be removed within 12 months of their formation.
However, a more recent study by Weller et al. (2004)
using data collected during the cooperative Subduction
Experiment from June 1991 to June 1993 indicated that
mode waters were formed and subducted during the
2-yr study. Weller et al. (2004) note that during this
period there appeared to be net heating over the region, in contrast to a net cooling suggested by the da
Silva et al. (1994) climatology. It is possible that subduction in this region is dependent on the climate state.
The warm conditions in the early 1990s lead to the
formation of a surface cap allowing the subduction to
occur (Weller et al. 2004); this is consistent with the role
of surface heating in the Lagrangian subduction process
proposed by Nurser and Marshall (1991). In the early
1980s the conditions may have been such that this cap
did not form and the mode waters were reentrained
into the mixed layer before being subducted.
The third type of mode water found in the North
Atlantic is a denser water mass formed on the southern
side of the North Atlantic Current (NAC). These are
the warmer range of the subpolar mode waters
(SPMW) described by McCartney (1982) and are identified as waters with a potential temperature between
10° and 15°C, salinity between 35.5 and 36.2 psu, and a
potential density in the range of 26.9–27.0. The ventilation of the eastern North Atlantic by SPMW suggested by McCartney (1982) and McCartney and Talley
(1982) was confirmed by Paillet and Arhan (1996b) using a set of hydrographic sections taken between 1977
and 1990, in conjunction with a large-scale thermocline
model. Paillet and Arhan (1996a) also showed the
equatorward subduction of 11°–12°C SPMW at (42°N,
12°W) on the eastern side of the basin using data collected along the Bord–Est hydrographic section between 20° and 60°N.
b. Mode water formation mechanisms
Tsujino and Yasuda (2004), in the opening paragraph
of their paper, give a very concise description of the
processes involved in the formation of subtropical
mode waters. The key features are 1) northward transport of warm waters from low latitudes into the subtropics by the western boundary current of the winddriven gyre; 2) a large exchange of heat between the
ocean and atmosphere in the subtropics, with a large
release of heat to the atmosphere during winter leading
to the formation of a deep surface mixed layer of ver-
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OLD AND HAINES
tically homogenized water; and 3) the transport of these
homogeneous waters by the southward interior Sverdrup flow across regions of shoaling winter mixed layer
depth into the permanent thermocline, where they are
insulated from the atmosphere, allowing them to persist
as a vertically homogenized layer (thermostad) or
mode water.
For a GCM to successfully generate mode waters, it
must include a realistic mixed layer model that allows
for both spatial and temporal variability in the mixed
layer depth. The importance of the mixed layer structure in setting the subtropical ventilation rate, and
hence allowing mode waters to form, was noted by
Woods (1985) and confirmed by Williams (1989). New
et al. (1995) showed that ventilation of the subtropical
gyre occurs from the southern side of a band of deep
winter mixing that stretches across the central North
Atlantic, which they called the subduction zone. Paillet
and Arhan (1996a) used a Lagrangian model to show
that subduction occurs along a line of zero net buoyancy input to the mixed layer. New et al. (2001) noted
that these two views are complementary as the line of
zero net buoyancy input corresponds to where shoaling
of the mixed layer toward the equator occurs, which in
turn corresponds to the subduction zone.
Sverdrup (1947) proposed that, away from the western boundary currents, the flow in the upper ocean
(⬍2000 m) is predominantly driven by the wind stress.
He went on to show that to first order the meridional
velocity can be defined as a function of the wind stress
curl. In the North Atlantic the subtropical gyre is
bounded to the south by the easterly trade winds and to
the north by the westerly midlatitude winds. Between
these two wind belts the wind stress curl is negative;
therefore the interior Sverdrup flow is equatorward. It
is this southward geostrophic transport that carries relatively deep mixed layer waters across the shoaling
mixed layer base resulting in subduction of water
masses. The lines of zero wind stress curl bound this
region. North of the northern line of zero wind stress
curl the interior flow is poleward, contributing to the
flow of the cyclonic subpolar gyre. Changes in the overlying wind field will result in shifts in position of the
lines of zero wind stress curl, which will in turn alter the
direction in which volume anomalies formed near the
zero lines propagate. This suggests that the northern
line of zero wind stress curl plays an important role in
determining the fate of volume anomalies formed along
the Gulf Stream extension and NAC.
c. Mode waters in HadCM3
To identify the mode waters formed in the North
Atlantic/Arctic Ocean region of HadCM3, a volume
1129
census of the waters in T–S space was made. Figure 1
shows the census based on the 100 yr of data, with the
volume in each T–S bin defined as a percentage of the
total volume. The water volumes have been sorted into
0.5°C by 0.1 psu T–S bins. A significant feature of the
T–S volume census for HadCM3, seen in Fig. 1, is the
diagonal banding. This is a consequence of the discrete
layering in the model, where the banding highlights the
low vertical resolution in the deeper model cells. The
deep cold waters make up the largest percentage of the
total volume; however there are three distinct modes
formed in the warmer classes: a weak mode at 18.5°C,
36.5 psu; a stronger mode at 16.5°C, 36.3 psu; and a
strong mode at 14.0°C, 36.1 psu. These three modes are
comparable to the modes observed in the North Atlantic. In the analysis that follows the three mode waters
will be identified by their central temperatures; these
are 18.5° (STMW), 16.5° (MMW), and 14.0°C (SPMW).
The four key fields contributing to the formation of
mode waters (winter mixed layer depth, surface heat
flux, P ⫺ E, and wind stress curl) have been calculated
from the 100 yr of HadCM3 data and are presented in
Fig. 2. The winter mixed layer base varies both spatially
and interannually; for the analysis presented, the 100-yr
local maximum mixed layer depth (Fig. 2a) has been
used to define the winter mixed layer base (WMLB).
The equatorward shoaling of the mixed layer is clearly
defined and extends across the central North Atlantic;
this corresponds to the New at al. (1995) subduction
zone. There is an eastward increase in WMLB depth to
the north of the subduction zone, which extends into
the eastern North Atlantic. The general spatial structure and depth of the HadCM3 winter mixed layer base
compare well with the recent mixed layer depth climatology produced by Kara et al. (2003). The climatology
shows the same deep trough extending across the basin
between 30° and 50°N, with depths varying from 250 m
in the west to over 400 m in the east. The deep mixed
layer extends into the subpolar region and the Labrador Sea. Over the subtropical gyre, the climatology
shows the mixed layer to be shallower than 150 m, as
does HadCM3, with the trough formed at 30°N on the
eastern side of the basin beneath the trade winds. The
mixed layer near the equator is shallower than 50 m in
both the climatology and HadCM3.
Figure 2b shows the 100-yr mean surface heat flux
into the ocean. There is a region of strong heat loss over
the Gulf Stream and its extension at the western end of
the deep WMLB trough. The depth of the WMLB in
this region is directly related to the significant heat loss
increasing the density and decreasing buoyancy. Equatorward of the subduction zone, on average there is
heating of the sea surface. This will tend to increase
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VOLUME 19
FIG. 1. A T–S volume census representing each class as a percentage of the total volume for
the North Atlantic region. To highlight the key mode water classes (⫹) formed in HadCM3
a log scale has been used to represent the volume percentages.
buoyancy leading to the shallower mixed layer depths
observed in Fig. 2a. The surface flux climatology, developed by Josey et al. (1998), has a region of high heat
loss, with an annual average of ⬎150 W m⫺2, over the
Gulf Stream, and regions of high heat loss around Iceland and in the Labrador Sea. These patterns and values are reproduced in HadCM3. The annual average
heating off Newfoundland is observed to be of the order of 50 W m⫺2, as produced by HadCM3. Over the
subtropical gyre there is an annual average heating of
around 25 W m⫺2, with a trough of heat loss of 25 W
m⫺2 beneath the trade winds. Hence the magnitude and
spatial pattern of surface fluxes produced by HadCM3
are consistent with the Southampton Oceanography
Centre (SOC) climatology.
The reverse of this pattern is seen in the 100-yr mean
P ⫺ E (Fig. 2c). The positive values (net precipitation)
poleward of the subduction zone correspond to freshening of the surface layer and hence increased buoyancy. Net evaporation occurs south of the subduction
zone leading to an increase in salinity and hence decreased buoyancy. Near the western coast of North Africa there is a region of strong net evaporation (related
to the trade winds) corresponding to a region of deep-
ening WMLB on the eastern side of the subtropical
North Atlantic. These general patterns are consistent
with the P ⫺ E maps produced by Schmitt et al. (1989)
based on the Bunker (1976) heat flux estimates and the
Dorman and Bourke (1981) precipitation estimates.
The observations show a peak in evaporation in the
eastern subtropics of 4.8 ⫻ 10⫺5 kg m⫺2 s⫺1, compared
with 5 ⫻ 10⫺5 kg m⫺2 s⫺1 in HadCM3. The observations
show a higher evaporation rate (⬎6 ⫻ 10⫺5 kg m⫺2 s⫺1)
over the Gulf Stream than produced in HadCM3 (5 ⫻
10⫺5 kg m⫺2 s⫺1). In contrast, HadCM3 shows a higher
rate of net precipitation (3.5 ⫻ 10⫺5 kg m⫺2 s⫺1) off
Newfoundland compared with the observations (1.5 ⫻
10⫺5 kg m⫺2 s⫺1). Also the line of zero net P ⫺ E is
farther south in HadCM3 than in the observations. The
difference in P ⫺ E over the Gulf Stream and its extension may result in local differences in the formation
processes.
The 100-yr mean wind stress curl is presented in Fig.
2d. The patterns of mean wind stress curl agree well
with the climatology produced by Harrison (1989). The
peaks in the wind stress curl climatology occur over the
Gulf Stream (8 ⫻ 10⫺8 Nm⫺3), in the eastern subtropics
(⫺10 ⫻ 10⫺8 Nm⫺3), off the west coast of North Africa
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OLD AND HAINES
FIG. 2. (a) The 100-yr max WMLB depth (m). The light shading highlights depths ⬎300 m, and the darker
shading depths ⬎500 m. (b) The 100-yr mean surface heat flux for the North Atlantic Region of HadCM3 (W m⫺2).
The shaded region highlights negative values. (c) The 100-yr mean P ⫺ E (⫻10⫺5 kg m⫺2 s⫺1). (Negative values
are shaded.) (d) The 100-yr mean wind stress curl (⫻10⫺8 N m⫺3). (Negative values are shaded.)
(8 ⫻ 10⫺8 Nm⫺3), and around Greenland (20 ⫻ 10⫺8
Nm⫺3). The magnitudes produced by HadCM3 compare well with the climatology. The lines of zero wind
stress curl define the separation between regions of
poleward and equatorward Sverdrup interior flow. The
northern line of zero wind stress curl slopes to the
northeast across the basin. All points along this line are
poleward of the shoaling of the WMLB. The meridional gradient in the wind stress curl on the western side
of the basin is steep, suggesting that there is little spatial
movement in the zero line. The meridional gradient is
weak on the eastern side of the basin, indicating that
the zero line may show greater variability in latitude,
which will affect both the formation and equatorward
propagation of the volume anomalies formed on the
eastern side of the basin. The line of zero wind stress
curl that bounds the equatorward side of the subtropical gyre region shows the reverse pattern of meridional
gradient, with steep gradients in the east and shallow
gradients in the west.
Marshall et al. (1993) estimated the annual kinematic
subduction rate using the equation
Sann ⫽ ⫺uH ⭈ ⵱H ⫺ wH.
共1兲
The overbar denotes annual means, H is the depth of
the winter mixed layer base (defined by the fixed
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FIG. 3. The 100-yr mean kinematic subduction rate (m yr⫺1)
calculated using Eq. (1). Contours in the clear areas give the
positive subduction rates, contours in the shaded areas give the
negative subduction rates (obduction), and the heavy contours are
the zero lines. The contour interval is 25 m yr⫺1.
WMLB surface), uH is the horizontal velocity at the
depth H of the WMLB, and wH is the vertical velocity
at the WMLB. Equation (1) was evaluated for the
North Atlantic region of HadCM3 using the numerical
method described by Marshall et al. (1999). The 100-yr
mean annual subduction rate for the North Atlantic
region of HadCM3 is presented in Fig. 3. Large subduction rates occur in the region of the subduction
zone, with a peak of approximately 100 m yr⫺1. This
value corresponds well with those calculated by Marshall et al. (1993) and Qiu and Huang (1995) using the
Levitus (1982) climatology for the North Atlantic. New
et al. (1995) also found similar subduction rates in their
modeling studies. This implies that HadCM3 is subducting water masses at realistic rates, and hence is
probably forming mode waters at a sensible rate.
3. Anomalous mode water formation and
subduction
a. Method of definition
Volume anomalies in temperature space will be
manifest as local thickness anomalies between pairs of
isothermal surfaces bounding the water mass in question. For the purpose of this analysis the three warmer
mode waters observed in HadCM3 are identified by
VOLUME 19
pairs of isothermal surfaces that bound the central temperature for each mode defined in Fig. 1. The STMW is
identified by the waters between 18° and 19°C, the
MMW by the waters between 16° and 17°C, and the
warm SPMW by the waters between 14° and 15°C.
Because of the discrete z levels used, the ocean
model does not have well-defined isothermal surfaces.
The heat content is represented in each fixed-volume
grid cell by a single potential temperature; consequently there is no resolved vertical distribution of temperature within the cells. Isothermal surfaces have been
defined from the discrete temperature field by linear
interpolation along the z coordinate between the midpoints of adjacent cells. This creates a smoothed representation of the vertical temperature profile and gives a
time series of isothermal depth that changes smoothly,
rather than in discrete steps defined by the vertical dimension of the grid cells. It should be noted that this
interpolation scheme does not conserve heat, in that
the full water column heat content calculated using the
interpolated temperature profile is in error by a few
percent. As only the spatial variability, rather than the
budget closure, is to be considered, this approximation
will suffice for the analysis presented.
Time series of the thickness between the pairs of
isothermal surfaces defining each mode were calculated
for every horizontal grid cell using the monthly mean
temperature fields taken from the 100-yr dataset. The
time series were filtered using a low-pass Lanczos filter
with a cutoff period of 3 yr to remove the seasonal
signal. From the filtered data, time series of the thickness anomalies relative to the 100-yr mean thickness
were calculated. These time series of thickness anomalies are used in this section to map the formation and
subduction regions for the modes considered, and in
section 4 to analyze the propagation of volume anomalies.
To identify the formation regions and subduction regions, the thickness anomalies were calculated for two
separate cases: (i) the total thickness anomaly at each
horizontal grid point, and (ii) the thickness anomaly at
each grid point for waters below the 100-yr maximum
WMLB. In case (i), the largest thickness anomaly will
occur where the isotherms outcrop in the wintertime
mixed layer. These correspond to the formation regions
of the mode waters. Hence the largest thickness variations in time will be due to the spatial movement of the
wintertime isothermal outcrops from year to year. This
variability is associated with changes in the formation
of the mode waters; therefore locations of maximum
variability for case (i) indicate regions of large anomalous surface formation.
Below the WMLB the mode water layers are wedge
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OLD AND HAINES
shaped, with the thin end of the wedge being equatorward of the subduction region. This thinning is a consequence of the conservation of potential vorticity
(PV), where
f⫹␵
.
PV ⫽
h
共2兲
Mode waters are formed with a low PV and away from
the western boundary current f Ⰷ␵. Therefore, as the
mode waters are advected equatorward away from
their formation regions, the planetary vorticity f decreases; therefore the thickness of the layer h must decrease to conserve PV. Hence for case (ii) the largest
thickness occurs where the isotherms intersect with the
WMLB. This corresponds to the regions where subduction or obduction of water volume will occur. These are
also the locations where the largest variability will occur; therefore locations of maximum variability for case
(ii) indicate the regions of large anomalous mode water
subduction.
b. Mode water formation regions
Spatial maps of the standard deviation in total thickness anomalies used to identify the formation regions
for the three modes are shown in Figs. 4a,c,e. All three
modes have broad regions of large variability showing
the extent of the formation regions. The STMW (18°–
19°C) spans 27.5° to 37.5°N and 70° to 45°W (Fig. 4a),
the MMW (16°–17°C) spans 30° to 41°N and 60° to
25°W (Fig. 4c), and the warm SPMW spans 35° to 45°N
and 50° to 10°W (Fig. 4e). The large spatial extent of
the formation regions is due to two possible processes.
1) The horizontal movement of the seasonal mixed
layer isothermal structure, which is equivalent to
changes in the wintertime sea surface temperature pattern. 2) The interannual changes in the depth of the
winter mixed layer base, which is related to the strength
of the wintertime deep convection. In practice both of
these processes are likely to occur simultaneously.
All three modes show a double peak in their formation region, suggesting a bimodal state during this 100yr period. The formation regions for the three modes
together span the west–east extent of the North Atlantic subduction zone, and there is minimal overlap between the three formation regions. This shows a shift in
mode water temperature to cooler classes from west to
east across the basin. The magnitude of the peak thickness anomaly increases from 70 m in the west (STMW)
to 100 m in the east (SPMW). However, we do not see
a continuous range of modes between the STMW and
SPMW (see Fig. 1), indicating that there are processes
1133
that inhibit the formation of waters in the intermediate
temperature ranges.
c. Mode water subduction regions
Spatial maps of the standard deviation in the thickness anomalies below the WMLB used to identify the
subduction regions for the three modes are shown in
Figs. 4b,d,f. The STMW (18°–19°C) subduction region
(Fig. 4b) forms a narrower band than the formation
region while retaining the corresponding double peak
structure. There is a weak peak at 28°N, 56°W and a
strong peak at 30°N, 49°W. A meridional section of
potential temperature taken at 50°W (Fig. 5a) for
March (i.e., the end of winter) of year 3 of the dataset
shows that 30°N corresponds to the intersection of the
18° and 19°C isothermal surfaces with the WMLB, suggesting that this is the latitude of subduction for the
STMW.
The subduction region for the MMW (16°–17°C; Fig.
4d) is broader than that of the STMW and shows two
very distinct peaks at 33°N, 39°W and 34°N, 26°W. A
meridional section of potential temperature taken at
30°W (Fig. 5b) for March of year 3 shows that 34°N
corresponds to the intersection of the 16° and 17°C
isothermal surfaces with the WMLB, suggesting that
this is the latitude of subduction for the MMW.
The main subduction region for the warm SPMW
(14°–15°C), shown in Fig. 4f, is very close to the eastern
boundary of the basin and in the region of deepening
WMLB (see Fig. 2a). There is a weak double peak
structure in the warm SPMW subduction, with a strong
peak at 36°N, 21°W and a weak peak at 39°N, 18°W.
Figure 5c shows the meridional section of potential
temperature taken at 20°W for March of year 3, highlighting the relationship between the isothermal surface
and the strong subduction peak at 36°N, 21°W.
d. Variability in mode water subduction
Time series of unfiltered thickness anomalies below
the WMLB at the main subduction locations (Fig. 5) for
each mode are presented in Fig. 6. The two variability
maxima observed for each mode will be identified as
the western peak (Fig. 6a) and the eastern peak (Fig.
6b).
The SPMW (14°–15°C) layer thickness (thick dashed
lines) shows large variability at both the western and
eastern locations during the first 50 yr of the dataset,
with some anticorrelation between them. Over the second 50 yr, there is very little thickness variability in the
east. The thickness at the western location varies
strongly over time scales of 5 to 20 yr, with the largest
variations occurring on the longer time scales. The
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FIG. 5. Three meridional sections taken from March of year 3 in the dataset showing the late
winter (Northern Hemisphere) potential temperature structure in the North Atlantic Ocean.
(a) Meridional section at 50°W highlighting the 18°–19°C STMW. (b) Meridional section at
30°W highlighting the 16°–17°C MMW. (c) Meridional section at 20°W highlighting the
14°–15°C SPMW. The dashed lines define the upper and lower bounds of the seasonal mixed
layer.
MMW (16°–17°C) thickness (thin solid lines) shows
large variability at both the western and eastern locations throughout the 100-yr period, and the two time
series also show some anticorrelation. There are periods when the thickness below the WMLB goes to zero
at the eastern location. This is due either to the movement of the isothermal outcrop southward, away from
this location, or the vertical movement of the layer so
that it intersects the fixed WMLB farther south. The
STMW (18°–19°C) thick solid lines show strong vari-
←
FIG. 4. Maps of std dev in the thickness anomaly (m) between pairs of isothermal surface in the total water column (formation) and
below the WMLB (subduction) for the three classes considered. (a) Formation and (b) subduction regions for the 18°–19°C water class.
(c) Formation and (d) subduction regions for the 16°–17°C water class. (e) Formation and (f) subduction regions for the 14°–15°C water
class. The heavy line shows the furthest extent of the isothermal outcrop. The contour interval is 10 m, and the location of the formation
and subduction peaks for each class is indicated.
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FIG. 6. Unfiltered time series of layer thickness (m) below the WMLB for the three mode
water classes taken from (a) the western and (b) the eastern peaks shown in Fig. 4. The
western peak locations are as follows: STMW (28°N, 56°W), MMW (33°N, 39°W), and warm
STMW (36°N, 21°W). The eastern peak locations are as follows: SPMW (30°N, 49°W), MMW
(34°N, 26°W), and warm SPMW (39°N, 18°W).
ability throughout the 100-yr period, with larger variability at the eastern location. The western location
shows some 20-yr periodicity, whereas at the eastern
location there are a series of separate strong events that
each last approximately 5 yr. The thickness of the
STMW in the eastern location periodically goes to zero
as the isothermal outcrop of this mode moves equatorward from this location.
Comparing the thickness variability between the different modes for the western locations, Fig. 6a shows
that, over the first 40 yr of the dataset, the events in the
modes appear to be lagged in time starting with the
STMW on the western side of the basin and ending with
the SPMW on the eastern side of the basin. The lag
between the STMW and SPMW events are of the order
of 8 yr. At present it is not clear what causes this lag.
The most probable candidates are the propagation of a
volume anomaly along the North Atlantic Current, a
gradual meridional shift in the wind field, and changes
in the surface heat fluxes. The complex interaction between these three processes makes it difficult to identify a leading cause. During the second half of the 100yr dataset, there appears to be little correlation between the modes.
Thickness variability events for the eastern locations
(Fig. 6b) for all modes are generally larger than those at
the western locations. Between years 15 and 30 there
are coherent events of the same sign in all three modes,
suggesting a basinwide change in the forcing. After 40
yr the SPMW (14°–15°C) shows very little variability,
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indicating that the main subduction variability occurs at
the western location. However, for both the STMW and
MMW there are periodic large events in the eastern
subduction locations.
These data indicate that formation and subduction
variations occur on periods of 5 to 20 yr. The dual
subduction locations appear to result from a change in
the gyre dynamics and/or a shift in the wind field leading to a northeast shift in the main subduction regions
for all modes. Changes in the SPMW formed on the
eastern side of the basin tend to lag changes in the
STMW formed on the western side by around 8 yr, as
seen in Fig. 6a by the offset in peaks between locations.
The northeast shift in the subduction region produces
larger subduction events at the eastern variability peak
for each mode water class. As a consequence these
events are likely to produce the strongest coherent signals for propagating anomalies, and will therefore be
used in the following section to study the persistence of
volume anomalies.
4. Persistence and propagation of subducted
anomalies
a. Mapping propagating thickness anomalies
In section 3 it was shown that the regions of largest
thickness variability, and hence large anomalous subduction, are localized. To show that the subducted volume anomalies persist and propagate beneath the
WMLB, the thickness anomalies at the peak subduction locations (Fig. 4) for each mode were lag correlated against the thickness anomalies at all other locations. For example, Fig. 7a(i) shows the maximum
thickness correlation detected for any lag time. A map
of the lag times for the maximum correlation, at all
locations where Fig. 7a(i) is statistically significant,
shows the spatial propagation of the volume anomaly
[e.g., Fig. 7a(ii)]. Lagged regressions of the standardized thickness anomalies are used to define the percentage of the remote thickness variability explained by the
anomalous subduction [e.g., Fig. 7a(iii)].
By lag correlating the thickness anomalies at the
peak subduction locations for each mode against the
thickness anomalies in adjacent water mass classes
(e.g., Figs. 7b,c), it is also possible to show that the
volume anomalies are transformed into adjacent classes
from the original water mass class as the anomaly
propagates. This is equivalent to determining the way
in which the heat associated with the anomaly diffuses
as the anomaly propagates around the system. In general if the anomaly transforms to warmer classes, then
it moves up through the water column, conversely if the
1137
anomaly transforms to colder classes then it moves
downward through the water column.
In the following, the results for each mode will be
presented separately. For each mode a set of maps
showing (i) the maximum correlation coefficients, (ii)
lag times, and (iii) percentage of remote variability explained is presented. Lagged correlations against the
water mass defining the mode, and against thermally
adjacent water masses, are shown.
b. Subtropical mode waters (18°–19°C)
The lagged correlations for the two STMW subduction locations show similar patterns, with those of the
eastern location (30°N, 49°W) giving stronger signals;
therefore only these results are presented. Figure 7
shows the maximum correlation of STMW (18°–19°C)
thickness anomalies below the WMLB at 30°N, 49°W
against thickness anomalies in 18°–19°C waters (Fig.
7a), thickness anomalies in 19°–20°C waters (Fig. 7b),
and thickness anomalies in 20°–21°C waters (Fig. 7c).
The statistically significant correlations for thickness
anomalies in the same class (18°–19°C) of waters are
everywhere positive [Fig. 7a(i)]. The lag times [Fig.
7a(ii)] are nearly all positive, that is, variability lags
variability at the peak subduction location, which is
consistent with the advection of anomalies away from
the subduction region equatorward and westward
around the gyre. The time scale for the propagation is
on the order of 5 yr. The regression map in [Fig. 7a(iii)]
shows that over the majority of the region of statistical
significance more than 60% of the remote variability
can be explained by the variability in the subduction
region, with the percentage explained decreasing with
distance from the source.
The adjacent warmer classes also both show regions
of statistically significant positive correlation. The lag
times for the 19°–20°C waters [Fig. 7b(ii)] show that the
earliest impact on this water layer lags the STMW subduction anomaly at (30°N, 49°W) by approximately 1
yr. The lag times extend out to 5 yr, moving south and
west, and 50%–75% of the remote variability is typically explained by the anomalous subduction of STMW
[Fig. 7b(iii)]. There is a small region of positively correlated thickness anomalies in the 20°–21°C waters
[Fig. 7c(i)] at the southwest extreme of the 18°–19°C
propagation region. The earliest impact on this water
class lags the STMW subduction anomaly by approximately 2 yr [Fig. 7c(ii)], and the subduction generally
accounts for less than 60% of the remote variability
[Fig. 7c(iii)]. These results are consistent with the transformation of STMW volume anomalies upward toward
warmer classes as a mechanism for mode water decay
within the thermocline.
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FIG. 7. Spatial maps of (i) max lag correlation, (ii) lag time at max for statistically significant correlations, and (iii) the percentage of
remote variability explained in thickness anomalies below the WMLB relative to the point (x) of max 18°–19°C subduction variability
at 30°N, 49°W, as defined in Fig. 4b. Correlated thickness variability of (a) 18°–19°C waters, (b) 19°–20°C waters, and (c) 20°–21°C
waters against variability in 18°–19°C subduction thickness. All data were low-pass filtered using a Lanczos filter with a 3-yr cutoff
period.
The transformation of STMW to warmer classes is
consistent with the transformation path that was found
to dominate in the basin-integrated diagnostics derived
by Haines and Old (2005). Using the integrated diagnostic it was possible to quantify the mixing processes
responsible for the transformations. Here it is possible
to see in detail the spatial distributions, range of lag
times, and the percentage of variance explained, and
hence the predictable variance in the different locations.
Lag correlations of STMW subduction anomalies
against thickness anomalies in the colder classes 17°–
18°C, 16°–17°C, and 15°–16°C are presented in Figs.
8a,b,c, respectively. For the 17°–18°C water mass (Fig.
8a), immediately beneath the STMW there is a dipole
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FIG. 8. Same as in Fig. 7, but at 30°N, 49°W. Correlated thickness variability of (a) 17°–18°C waters, (b) 16°–17°C waters, and (c)
15°–16°C waters against variability in 18°–19°C subduction thickness.
pattern of correlation [Fig. 8a(i)]. To the west of the
subduction region the anomalies are negatively correlated, while to the east they are positively correlated.
The lag times [Fig. 8a(ii)] again suggest the propagation
of anomalies southward around the gyre away from the
subduction region, and for both regions the lag times
are up to 5 yr. For the western negative correlation
region, up to 80% of the remote variance is explained
by the subduction anomalies, while in the eastern
positive correlation region, less than 60% is explained
[Fig. 8a(iii)].
The positively correlated region to the east cannot be
due to propagation from the subduction region as there
is a zero lag for the earliest impact; this suggests that
this correlation results from the correlated formation of
waters in the 17°–18°C water mass. This is consistent
with the observed decrease in dominant mode water
temperature moving eastward across the basin. The an-
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ticorrelated region to the west suggests that the variability here results from the anomalous formation of
the 18°–19°C waters at the expense of 17°–18°C waters
through a shift in the formation/subduction regions.
This interpretation is also consistent with the extent of
the high thickness variability for the 18°–19°C mode
waters in Fig. 4a, which extends westward from the
peak at 30°N, 49°W.
The two colder classes (16°–17°C and 15°–16°C)
show similar patterns of anticorrelation [Figs. 8b(i) and
8c(i)], with lag times out to 5 yr [Figs. 8b(ii) and 8c(ii)].
For the 16°–17°C class, up to 80% of the remote variability is explained by the anomalous subduction [Fig.
8b(iii)], while up to 70% is explained for the 15°–16°C
waters [Fig. 8c(iii)]. This pattern of anticorrelation is
repeated down to 13°C waters, below which there is no
significant correlation (data not presented). This suggests that subduction of a positive anomaly of 18°–19°C
water is associated with a loss of waters in the lower
thermocline classes beneath the subduction region.
If the thickness of the 18°–19°C water class below the
fixed WMLB is increased, then the thickness of some
other temperature layers must decrease at the same
time to make room for the 18°–19°C water. However,
the integrated transformation diagnostics in Haines and
Old (2005) do not indicate any transformations between the 15°–16°C and the 18°–19°C water classes.
This implies that the reduction in thickness of the
colder waters is a dynamical effect, with divergence/
convergence of the colder classes allowing for the presence of the 18°–19°C mode water anomalies above.
c. Madeira Mode Waters (16°–17°C)
The Madeira Mode Waters (16°–17°C) also show two
distinct locations of maximum subduction variability.
Lag correlation results for both peaks show different
responses; therefore, the analysis for both locations will
be presented. The analysis for the western subduction
location at 33°N, 39°W is shown in Fig. 9, and the analysis for the eastern location at 34°N, 26°W in Fig. 10.
Spatial maps of the correlations of the MMW subduction anomalies against the 15°–16°C, 16°–17°C
(MMW), and 17°–18°C thickness anomalies are presented for the two subduction regions.
The main results for the western peak at 33°N, 39°W
are that the subducted anomalies propagate away from
the subduction region around the subtropical gyre
[Figs. 9b(i) and 9b(ii)] with significant correlations seen
for up to 8 yr. This is considerably longer than for the
STMW, which is perhaps consistent with the location of
the MMW subduction, that is, farther from the center
of the gyre and in a region of lower flow. Unlike the
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STMW, the MMW thickness anomalies are anticorrelated with the anomalies in the layer immediately above
them in the water column [17°–18°C; Fig. 9c(i)] and
these anticorrelated anomalies do not appear to propagate very far around the gyre. However, for the colder
water class (15°–16°C) there is a double branch in the
correlation pattern, with both branches being positively
correlated with the subduction anomalies [Fig. 9a(i)].
The lag times [Fig. 9a(ii)] in both branches extend out
to 7 yr. Both branches start with zero lag in relation to
the subduction anomaly, suggesting that they result
from the correlated formation of waters in this colder
class. This correlated formation will mask any signal of
the transformation of waters with adjacent classes. This
observation suggests that the anomalies are not removed via transformation into adjacent classes; an alternative mechanism for their removal will be discussed
in section 5.
The eastern peak at 34°N, 26°W produces larger
anomalies and longer lag times [of the order of 10 yr in
Fig. 10b(ii)]. The correlation of the subduction anomalies against the 16°–17°C waters [Fig. 10b(i)] shows a
double branch, with the western branch being anticorrelated and the eastern branch positively correlated.
The western branch coincides with the region of positive correlation for the western subduction peak at
33°N, 39°W seen in Fig. 9b(ii). This anticorrelation suggests that the dual subduction peaks result from a shift
eastward in the isotherm structure. Figure 10a(i) shows
the thickness anomalies in the waters beneath the
MMW to be anticorrelated with the anomalies in the
MMW subduction region, that is, the colder classes are
displaced by the subduction of positive volume anomalies. The anomalies in the waters immediately above
the MMW [Fig. 10c(i)] are positively correlated with
the subducted anomalies. The shortest lag time is 0 yr
immediately adjacent to the subduction region and extends out to 10 yr at the southern limit of the statistically significant region. The anomalies in this layer
travel faster than those in the MMW layer, suggesting
that the observed correlations are more likely to be due
to coformation rather than transformation between
classes.
d. Warm subpolar mode waters (14°–15°C)
For the warm SPMW class the western subduction
peak at 36°N, 21°W shows strong correlation signals,
while the weaker eastern peak at 39°N, 18°W does not;
therefore, only the results for the western peak are
shown in Fig. 11. The figure includes maps of the correlation of the warm SPMW thickness anomalies at the
subduction location against the 13°–14°C (Fig. 11a),
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1141
FIG. 9. Same as in Fig. 7, but for 16°–17°C subduction variability at 33°N, 39°W, as defined in Fig. 4d. Correlated thickness variability of
(a) 15°–16°C waters, (b) 16°–17°C waters, and (c) 17°–18°C waters against variability in 16°–17°C subduction thickness.
14°–15°C (warm SPMW; Fig. 11b), and 15°–16°C (Fig.
11c) thickness anomalies.
There is an extensive region of strong positive correlation [Fig. 11b(i)] associated with the thickness
anomalies in the subduction region. The lag times in
Fig. 11b(ii) show a broad area of essentially zero lag
around the subduction location, highlighting the wide
extent of the subduction region. Once again lag times of
up to 10 yr imply the propagation of the anomalies
away from the subduction region equatorward and
westward around the gyre. Over most of the region,
more than 80% of the remote variability is explained by
the thickness anomalies in the subduction region [Fig.
11b(iii)], that is, the subducted volume anomalies of
warm SPMW produce propagating signals that are coherent for more than 10 yr and hence should be predictable. This strong coherence also means that heat
anomalies produced by anomalous subduction in this
class will persist for up to 10 yr.
The warmer water mass (15°–16°C) immediately
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FIG. 10. Same as in Fig. 9, but at 34°N, 26°W.
above the warm SPMW has a region of negative correlation [Fig. 11c(i)] above the SPMW subduction region. This can be interpreted as a shift in the isothermal
outcrops equatorward, leading to the formation of
cooler 14°–15°C waters at this location. Equatorward of
the SPMW subduction region there is an extended region of positive correlation. The lag times [Fig. 11c(ii)]
and regression coefficients [Fig. 11c(iii)] indicate that
this region of positive correlation in the warmer class is
due to correlated formation. This is also consistent with
the equatorward movement of the isotherms leading to
a change in position of the formation regions for waters
masses that are adjacent in temperature class. Figure 5c
also supports this interpretation, as it shows that the
shoaling region of the WMLB has a shallower slope
near the eastern side of the basin. There are thick layers
of both 14°–15°C and 15°–16°C waters beneath the
WMLB along this meridional section at 20°W. The correlations with the 16°–17°C waters (data not shown) are
negative, indicating the displacement of these warmer
waters by the SPMW classes below.
In contrast, the thickness anomalies in the colder
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1143
FIG. 11. Same as in Fig. 7, but for 14°–15°C subduction variability at 36°N, 21°W, as defined in Fig. 4f. Correlated thickness variability of
(a) 13°–14°C waters, (b) 14°–15°C waters, and (c) 15°–16°C waters against variability in 14°–15°C subduction thickness.
class of waters (13°–14°C), shown in Fig. 11a, are positively correlated with the subducted thickness anomalies of the warm SPMW (14°–15°C). Again the lag times
and regression coefficients suggest that this correlation
is the result of correlated formation. It is interesting to
note that the propagation of the signal in this colder
class is interrupted around 25°N. This corresponds to
the region of deepening WMLB (as seen in Fig. 2a) due
to the increased evaporation by the trade winds. Figure
5a also shows that there is a strong 18°–19°C ther-
mocline formed in this region that tends to pinch off the
colder classes below. These features suggest that this is
an obduction region (consistent with Fig. 3), where the
lateral advection is carrying waters back into the seasonal mixed layer across the WMLB.
5. Discussion
It has been shown that persistent volume anomalies
form in the warm mode waters of the North Atlantic
Ocean in the HadCM3 climate model. Analysis of the
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100 yr of data taken from the control run has shown
that the time scales associated with the variability in
formation of North Atlantic mode water ranges from 5
to 20 yr. For some events there was a west–east lag of
8 yr in the anomalous formation across the basin. This
is consistent with the observed time scale for the propagation of heat anomalies in the upper ocean along the
North Atlantic Current (Hansen and Bezdek 1996; Sutton and Allen 1997), suggesting that this is one possible
cause for the lagged anomalous formation across the
basin.
The key process leading to the persistence of these
anomalies is their subduction beneath the winter mixed
layer base where they are effectively isolated from the
surface forcing. Once beneath the WMLB, the volume
anomalies propagate equatorward away from the subduction region and move as part of the general gyre
circulation. The persistence time of volume anomalies
increases with distance east of the center of the subtropical gyre. This is consistent with the increase in
subduction depth and decrease in circulation with distance east of the subtropical gyre center. The STMW
anomalies formed on the western side of the basin typically persist for up to 5 yr after they are subducted
beneath the winter mixed layer base. Moving eastward,
MMWs are formed over a region spanning from 40°W
above the mid-Atlantic ridge to 20°W. The subducted
MMW anomalies persist for up to 8–10 yr as they travel
around the gyre in the general circulation. On the eastern side of the basin the SPMW anomalies subducted
beneath the WMLB persist for up to 10 yr as they
propagate around the gyre.
For all three modes it was shown that anomalous
volumes are on average formed at the expense of the
colder temperature classes; therefore, the formation of
a positive volume anomaly results in warming, or ocean
heat storage. Because of the shorter propagation times,
and hence shorter persistence, nearer the gyre center,
the STMW will only contribute to ocean heat anomalies
on time scales of up to 5 yr. The persistence times of up
to 10 yr for the modes formed on the eastern side of the
basin, implies that they play an important role in the
decadal storage of ocean heat anomalies. The persistence of volume (heat) anomalies is equivalent to the
ocean’s memory of warming or cooling climatic events.
The persistence times of volume anomalies in particular temperature ranges can be seen in a time series
of the basin-integrated volumes based on the potential
temperature. Figure 12 presents Hovmöller plots of the
basin-integrated volumes between pairs of isothermal
surface that are 0.5°C apart in temperature space. The
data have been detrended and low-pass filtered, using a
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Lanczos filter with a cutoff period of 1.5 yr, to remove
the seasonal cycle.
Figure 12a presents a Hovmöller plot of the time
variation in the total volumes of water above the
WMLB, sorted into 0.5°C temperature bands. This figure shows the formation of volume anomalies around
18°C and the subsequent propagation to colder classes
in time. As these volumes are above the WMLB, this
process occurs completely within the seasonal mixed
layer. As these upper-ocean heat anomalies propagate
across the North Atlantic basin along the Gulf Stream
extension/NAC, they lose heat to the atmosphere and
are transformed to colder water classes. For example, a
large volume anomaly forms in the mixed layer between 17° and 18°C around year 32 and subsequently
propagates to colder classes over the following 8 yr.
Figure 12b shows the volumes below the WMLB for
temperature classes in the range of 5° to 25°C. To the
right of the figure is a plot of the detrended, standardized, and low-pass-filtered (Lanczos, 1.5-yr cutoff) total
heat content anomaly for the 100 yr. This figure shows
that large volume changes occur in the classes around
the main North Atlantic Mode Waters and below the
WMLB. Around 40 yr into the time series there is a
large positive heat anomaly in the North Atlantic basin.
This is distributed amongst the main mode temperature
classes seen in Fig. 1, of 7°, 10°, 14° and 16°C. Figure
12b highlights the role that mode waters, particularly
those formed on the eastern side of the basin, play in
the long-term ocean heat storage.
To further determine whether particular classes have
greater variability than others, the rms of the anomalies
in time for each temperature class was calculated. In
general there is an increase in the magnitude of the
anomalies going from warmer to colder classes. This is
a consequence of the increasing contribution to the total volume of each class going from warm to cold
classes. To minimize this volume effect in the rms of the
anomalies, the rms values were weighted according to
the percentage their temperature class contributed to
the total basin volume. The weighted rms values, presented in Fig. 12c, show distinct peaks associated with
the main mode waters, in particular 14°–15°C, 16°–
17°C, and 18°C. It also highlights the colder subpolar
modes at 13°–14°C, 9°–10°C, and 6°C.
In the absence of any long-term changes in the total
heat content, any volume anomalies formed in a particular temperature range must be removed. For the
STMW it was shown that there is transformation of the
18°–19°C waters to warmer classes as the volume
anomaly propagates around the gyre. At the western
extent of the propagation region the volume anomalies
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1145
FIG. 12. Hovmöller plots of the detrended and low-pass-filtered basin-integrated volumes
between pairs of isothermal surfaces separated by 0.5°C in temperature space. The waters
have been separated into (a) volumes (⫻1014 m3) above the 100-yr max WMLB and (b)
volumes (⫻1014 m3) below the WMLB. For comparison the detrended, standardized, and
low-pass-filtered total North Atlantic/Arctic Ocean heat content [H(T )⬘] has been included to
the right of the Hovmöller plots. The data were filtered using a Lanczos low-pass filter with
a 1.5-yr cutoff period to remove the seasonal signal. (c) The weighted rms volume anomalies
below the WMLB in each 0.5°C temperature class. The weighting applied is the inverse of the
percentage of the total volume below the WMLB that the class mean volume represents.
are moving northward in the gyre recirculation; this
carries them across the shoaling region of WMLB back
into the mixed layer (i.e., obduction through horizontal
advection) and into a region of large heat loss to the
atmosphere. It is through the reemergence of the vol-
ume anomaly into the mixed layer that the ocean heat
anomaly can be removed through ocean–atmosphere
interaction. This, however, does not appear to be the
case for the volume anomalies formed on the eastern
side of the basin, that is, the MMW and SPMW, as these
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anomalies do not propagate far enough around the gyre
and are too deep to be obducted back into the mixed
layer.
For all three modes there are no coherent signals
associated with the subducted thickness anomalies
south of 20°N. There is a region of net heat loss from
the ocean to the atmosphere south of 20°N (Fig. 2b)
that lies beneath the trade winds. The trade winds form
the boundary of the tropical–subtropical water mass
exchange region discussed by Liu et al. (1994). They
propose that tropical–subtropical water mass exchange
is a consequence of the shallow (300 m) overturning cell
extending from the equator into the subtropics, driven
by the Ekman transport of the easterly trade winds.
This cell interacts with the equatorial cell that drives
the countercurrent. The pathway taken by a water parcel depends on the depth at which it reaches the Ekman
cell. Figure 5b shows that the MMWs are blocked by
the equatorial countercurrent (centered at 10°N) that
brings colder waters up from the western boundary to
the surface. Therefore, MMW anomalies will probably
rejoin the system at the western boundary near the start
of the Gulf Stream. Figure 5c indicates that the SPMW
feeds directly into the countercurrent region. Therefore
SPMW anomalies may return to the surface via the
equatorial countercurrent or by upwelling at the equator as part of the Ekman overturning cell.
The analysis of the HadCM3 data indicates that the
anomalies formed in the MMW and SPMW are not
significantly transformed to the surrounding classes as
they propagate around the subtropical gyre. The tropical–subtropical exchange region seems to be the most
probable region where the heat anomalies stored in the
deeper modes will reemerge to release their heat to the
atmosphere. It should be noted that Harper (2000)
showed, using tracer experiments, that this process is
complicated by the basin topography in the North Atlantic, and that it is possible that waters entrained into
this cell from the subtropics actually reemerge in the
Gulf of Mexico. More work is required to fully understand the reemergence of the deeper heat anomalies
and their impact on climate variability.
The long coherence time of the propagating volume
anomalies observed near the eastern side of the basin
implies the presence of signals in the ocean hydrography that may give predictability out to 10 yr. It should
be noted that these results are based on observations
from a fairly low resolution model. The ocean component of the model is not eddy permitting and will therefore not contain the high-frequency natural variability
that exists in the real world. This may affect the mixing
rates in the formation regions and will lead to a more
stable Gulf Stream extension, that is, it will preclude
VOLUME 19
meanders. The absence of these processes means that
the model signals are possibly too smooth, contributing
to the spatially coherent high correlation and regression
values observed in the analysis. In reality it may be
more difficult to find these signals in the observational
record. However, the model observations provide good
physical grounds for seeking persistent signals in this
part of the North Atlantic. The potential for volume
anomalies to enter the Tropics through tropical–
subtropical water mass exchange processes beneath the
trade winds suggests the possibility of a strong atmospheric feedback, and the time scales involved suggest a
potential climate feedback.
On much longer time scales, changes in the meridional overturning circulation (MOC) are likely to drive
large-scale climate changes through ocean basin exchanges. Density-driven processes (e.g., deep convective mixing/overturning) in the high latitudes of the Atlantic Ocean and the Arctic Ocean will influence, and
be affected by, the variability of the MOC. Figure 12
indicates that the colder SPMW (10°–11°C) and the
Labrador Sea Waters (6°–7°C) play an important role
in long-term ocean heat storage. [It should be noted
that Labrador Sea Waters in HadCM3 are warmer than
in the real word, as discussed by Cooper and Gordon
(2002).] The difficulty with these classes is that they are
strongly affected by changes in salinity (i.e., densitydriven processes), which is not resolved using this temperature class analysis. Hence it is very difficult to find
coherent correlations in these water masses. Coherent
signals in the colder classes show up more clearly in
potential density coordinates. A preliminary analysis
based on a volume census in potential density space
indicates that there are propagating signals in the
colder mode water classes similar to those for the
warmer classes described in this paper.
Acknowledgments. This work was supported by
NERC under the COAPEC thematic program (NER/
T/S/2000/00307).
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