OSOcean ScienceOSOcean Sci.1812-0792Copernicus PublicationsGöttingen, Germany10.5194/os-12-481-2016Compensation between meridional flow components of the Atlantic MOC at 26∘ NFrajka-WilliamsE.e.frajka-williams@noc.soton.ac.ukhttps://orcid.org/0000-0001-8773-7838MeinenC. S.https://orcid.org/0000-0002-8846-6002JohnsW. E.SmeedD. A.https://orcid.org/0000-0003-1740-1778DuchezA.LawrenceA. J.CuthbertsonD. A.McCarthyG. D.https://orcid.org/0000-0002-2363-0561BrydenH. L.BaringerM. O.MoatB. I.RaynerD.Ocean and Earth Science, University of Southampton, National Oceanography Centre Southampton, SO14 3ZH, UKAtlantic Oceanographic and Meteorological Laboratory, Physical Oceanography Division, 4301 Rickenbacker Causeway, Miami, FL 33149, USAUniversity of Miami, Rosentiel School of Marine and Atmospheric Science, 4600 Rickenbacker Causeway, Miami, FL, USANational Oceanography Centre, University of Southampton Waterfront Campus, European Way, Southampton, SO14 3ZH, UKE. Frajka-Williams (e.frajka-williams@noc.soton.ac.uk)1April20161224814939October201513November20158March201617March2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://os.copernicus.org/articles/12/481/2016/os-12-481-2016.htmlThe full text article is available as a PDF file from https://os.copernicus.org/articles/12/481/2016/os-12-481-2016.pdf
From ten years of observations of the Atlantic meridional
overturning circulation (MOC) at 26∘ N (2004–2014), we revisit the
question of flow compensation between components of the circulation.
Contrasting with early results from the observations, transport variations of
the Florida Current (FC) and upper mid-ocean (UMO) transports (top
1000 m east of the Bahamas) are now found to compensate on sub-annual
timescales. The observed compensation between the FC and UMO transports is
associated with horizontal circulation and means that this part of the
correlated variability does not project onto the MOC. A deep baroclinic
response to wind-forcing (Ekman transport) is also found in the lower North
Atlantic Deep Water (LNADW; 3000–5000 m) transport. In contrast,
co-variability between Ekman and the LNADW transports does contribute to
overturning. On longer timescales, the southward UMO transport has continued
to strengthen, resulting in a continued decline of the MOC. Most of this
interannual variability of the MOC can be traced to changes in isopycnal
displacements on the western boundary, within the top 1000 m and
below 2000 m. Substantial trends are observed in isopycnal
displacements in the deep ocean, underscoring the importance of deep boundary
measurements to capture the variability of the Atlantic MOC.
Introduction
The Atlantic meridional overturning circulation (MOC) is a key part of
the global ocean circulation, redistributing heat and properties around
the globe. The continuous daily time-series observations at
26∘ N (Fig. 1) are the first of their kind, capturing the transbasin
circulation variability on timescales of days to – now – a decade.
From the first year of observations, noted
considerable high-frequency variability of the MOC, with values ranging from
5 to 35 Sv. This exceeds the range of all previous estimates of the MOC
strength at 26∘ N derived from hydrographic sections
. After 3 years of data were recovered,
a seasonal cycle of the MOC became apparent , with
a maximum in July–November and a minimum in March, and a seasonal range of
approximately ±3.5Sv. This seasonal cycle has been captured by
numerical simulations and may be explained by variations
in wind-forcing on seasonal timescales .
also noted that the components of the MOC (the
Florida Current, the interior thermal-wind contribution, and the Ekman flow)
were largely uncorrelated, suggesting that each contributes variability to
the MOC independently. Once the MOC records had stretched to 7 years in
length, striking interannual variability and more recently a declining trend
in the MOC were revealed . These
low-frequency transport changes have been shown to be responsible for
changes in ocean heat content in the subtropical and tropical North Atlantic
.
Numerical investigations into the sources of variability to the Atlantic
MOC interannual variability suggest that much of the variability may be
attributable to winds
.
Buoyancy forcing instead affects decadal variations
. An estimate of the MOC from high latitude
density anomalies suggests a decline of the MOC ,
which is presently observed at 26∘ N (observed trend of
-0.5 Svyr-1; ) though it may not be
indicative of a longer-term decline . The
26∘ N array provides an estimate of the MOC, but separating it into
components and even depth ranges of anomalies may aid in the identification
of physical causes of change .
Map of the study area. The RAPID array is shown with dashed lines
crossing the Atlantic around 26∘ N. Mooring positions are given by
red squares. The inset is marked by the black rectangle, and shows a zoomed
in view of the western boundary region.
In this paper, we introduce the 10-year record of the MOC at
26∘ N, describing features of the variability in the most recent
18 months and across the 10-year record, and examine more fully the
degree of correlation or compensation between MOC components using the longer
records. While much of the recent research into Atlantic MOC variability has
focused on interannual timescales and longer, here we quantify newly observed
compensation between the Florida Current and UMO transports, and
co-variability between the deep transbasin transports and zonal winds, on
sub-annual timescales. The depth structure and timescales of these
variations are explored, illustrating an important role for the western
boundary below 1000 m. Lower-frequency changes in MOC components,
including the continuing trend in the vertical shear of the mid-ocean
transport, are also described. Finally, we conclude by discussing the origins
of the lower-frequency variability in the 10-year records.
Methods
The international 26∘ N Rapid Climate Change
(RAPID)/Meridional Overturning Circulation and Heat Flux Array (MOCHA), hereafter RAPID
26∘ N, has provided comprehensive daily measurements of the MOC at
26∘ N for 10 years (April 2004–March 2014;
). The MOC is defined as the northward transport
above the depth of maximum overturning (roughly 1100 m) across
26∘ N, and is constructed as the sum of three components: the
surface meridional Ekman transport estimated from reanalysis winds, the
Gulf Stream transport through the Florida Straits – the Florida Current (FC)
– measured by a submarine cable (e.g. ), and the
upper mid-ocean (UMO) transport, measured by a transbasin array of current
metre and dynamic height moorings between the Bahamas and Africa. The exact
number of moorings and instruments has varied over the past decade during
which there have been over 20 deployment and recovery cruises. The main
western boundary mooring that we used here is called WB2, and typically has
18 MicroCAT (Seabird Electronics, Bellevue, WA)
conductivity–temperature–depth instruments. Vertical resolution ranges from
75 m near the surface to 500 m near the bottom. Overall,
the accuracy of the MOC transport is estimated to be 1.5 Sv (10-day
values) or 0.9 Sv (annual averages). Full details of the array
configuration and map (their Fig. 1.1), transport calculation, and
associated errors can be found in .
Here, we focus on the transbasin or mid-ocean (MO) transport, from
which the UMO is derived. The MO transport is constructed from three
parts:
MO(z)=Twbw(z)+Tint(z)+Text(z),
where Twbw is the western boundary transport estimated from
direct current metre measurements, Tint the “internal”
transbasin transport, and Text the “external” flow. The western
boundary wedge transport, Twbw, includes most of the flow
associated with the Antilles Current. The internal transport,
Tint, is the baroclinic flow zonally integrated across the
remainder of the ocean interior relative to an assumed level of no motion at
4820 dbar. It is derived from dynamic height moorings near the western and
eastern boundaries and over the Mid-Atlantic Ridge (Fig. 1). Here we focus on
the western and eastern profile contributions to the Tint. Using
only the western and eastern density contributions to interior
transport per unit depth, Tint(z) relative to the
reference level (zref) is related to density as
Tint(z)=-gfρ∫zrefzρe(z′)-ρw(z′)dz′,
where g is gravitational acceleration, f the Coriolis parameter, and
ρe(z) and ρw(z) the density profiles at the
eastern and western boundary, respectively .
The external flow, Text, is the (unmeasured) interior
barotropic flow that ensures zero mass transport across the section. This
component is calculated as a residual of the other components and is applied
as a uniformly distributed, and thus depth-independent, velocity across the
entire mid-ocean section, which we refer to as hypsometric compensation.
Due to changes in the width of the basin as a function of depth, even though
the applied flow is barotropic, the transport per unit depth has decreasing
magnitude with increasing depth. showed that this
estimate – derived from mass conservation – was in good agreement with an
independent estimate of the mid-ocean barotropic flow derived from bottom
pressure gauges deployed across the section over the April 2004–April 2005
period.
The MO transport can be further divided into its contributions to the
upper and lower branches of the overturning circulation. The UMO is
defined as the depth integral of MO transport between the surface and
the time-varying depth of maximum overturning, roughly 1100 m. The lower
limb of the MOC is made up of southward flowing North Atlantic Deep
Water, which is split into contributions associated with upper North
Atlantic Deep Water (UNADW; 1100–3000 m) and lower North Atlantic Deep
Water (LNADW; 3000–5000 m). The sum of these two transports recovers
nearly all the variability of the MOC (r=0.996; ). The small difference is equal to the flow between
1100 m and the depth of maximum overturning and a contribution from the hypsometric compensation below 4820 bar.
For the analysis presented here, we start with the RAPID data as processed
for the publicly available data set. This processing involves filtering
individual instrument records with a 2-day low-pass filter to remove the
tides, and subsampling onto 12-hourly intervals. From this subsampled
data set, transport components are computed, then further 10-day low-pass
filtered with a fifth-order Butterworth filter before the compensation
transport is calculated . These data are available
from http://www.rapid.ac.uk/. Here we
additionally bin the data onto a twice-monthly time grid, then remove the
twice-monthly climatology to reduce seasonal variations. For lower-frequency
variations, deseasonalized time series are further filtered with
a 1.5 year Tukey filter. Significance and confidence intervals are
reported at the 95 % level, unless otherwise indicated. The number
of degrees of freedom was calculated using the integral timescale of
decorrelation to the first zero crossing . When
a year is denoted 2009/10, it refers to the period 1 April 2009 through
31 March 2010.
For the purpose of calculating isopycnal displacements ζ, absolute
salinities and conservative temperatures on the twice-monthly time grid are
used. Isopycnal (σ) displacements are then calculated following
, as
ζ(z¯(σi),t)=z(σi,t)-z¯(σi),
using locally referenced densities. For example, to determine the
displacement of the isopycnal typically found at 1000 dbar, potential
densities are calculated referenced to 1000 dbar (σ1). Using these
locally referenced densities, the mean density at 1000 dbar is identified
(σi=〈σ1(z=1000,t)〉t). The depth of this density,
z(σi,t), is then determined for each time step in the
locally referenced densities, and differenced from its mean depth
(∼ 991 m for 1000 dbar). This produces isopycnal displacements as a
function of density, which can then be mapped back onto depth using the mean
relationship between depth and density (z¯(σi)). This process of
locally referencing densities is repeated for each pressure surface from the
surface to the bottom at 20 dbar intervals. The detrended and
low-pass filtered time series are processed as above.
10-years of MOC and mid-ocean variability
All of the 10-year time series of transport components at 26∘ N
show high-frequency variability (Fig. 2a). In the most recent
18 months, additional features of the time series include a large
Ekman transport reversal in March 2013 (similar to the two reversals that
occurred in 2009/10 and 2010/11). During the March 2013 event, the Ekman
transport anomalies exceeded 2 standard deviations from the mean, with the
typically northward-flowing water turned to the south. This reversal was
similar in magnitude to the December 2009–March 2010 event, but with shorter
duration (Fig. 2a). On several occasions, as during the negative Ekman events
in 2005, 2010, and 2013, the FC also showed sharp, short-term reductions in
transport. These corresponding anomalies led to sharp reductions – or even
brief reversals – of the MOC at these times. Over the past 10 years, the
MOC was negative from 19 to 24 December 2009, and from 9 to 13 March 2013. The
Ekman transport reversals also coincided with reductions of the southward
LNADW flow (Fig. 2b). In the most recent 5 years, the LNADW
experienced more short periods of reversal (i.e. a northward flow of the net
transport below 3000 m) than had been observed in the first
5 years of the record. These high-frequency events in the deep flow
exhibit fairly weak vertical shear, with maximum anomalies below
3000 m (Fig. 3a).
(a) Transport time series of the FC (blue), Ekman (green),
upper mid-ocean (magenta), and overturning (red) at 10-day resolution.
(b) Layer transports for UNADW (1100–3000 m, cyan) and
LNADW (3000–5000 m, purple). For visualization purposes, the
filtered versions of the time series are shown in black, where transports
have been convolved with a 4-month, low-pass Hanning window. Transports are
positive northwards; 10-year mean transports for each components are shown
with the dotted lines.
Transport-per-unit-depth anomalies of the mid-ocean transport at
26∘ N, where the time-mean profile over the 10 years has been
removed. (a) The top panel shows the twice-monthly, deseasonalized
variability and (b) the lower panel is further filtered with the
1.5 year filter. Red (blue) shows transports that are anomalously
northward (southward).
Time series of transport anomaly for MOC and components as in
Fig. 2, deseasonalized and low-pass filtered with 1.5 year filter.
The MOC exhibits substantial variability at timescales longer than
annual (Fig. 4). Interannual variability of the MOC derives primarily
from the UMO component, with a negative trend in both over the full
10-year record. The mean and standard deviation of the MOC for the first
5 years of observations is 18.4±1.3Sv and the second 5
years 15.5±1.9Sv (these values are significantly different; see
Table 1). While some of this change is contributed by the 2009/10 dip , the intensification of the southward
thermocline flow (UMO) has persisted with the associated weakening of
the MOC. The 2012/13 year was the second weakest year of
the MOC (14.2 Sv), behind the 2009/10 year (12.8 Sv). In contrast to the
2009/10 year, the weak MOC in 2012/13 had very little contribution from
the wind-driven Ekman transport, but instead is associated with a strong
southward thermocline flow (UMO).
Mean ± standard deviation of transports for the first 5
years and latter 5 years, where years run from 1 April–31 March. Standard
deviations are calculated on the annual averages. Statistically significant
changes to the mean are indicated by bold, based on two-tailed t tests.
Transport-per-unit depth anomaly profiles show the depth structure of
mid-ocean transport variations. In the top 1100 m, the southward UMO has
intensified (Fig. 3b, shift from red to blue), while below 3000 m the
southward LNADW has weakened (shift from blue to red). Previous analyses
have shown that variability of the UMO on interannual timescales is
primarily governed by changes at the western boundary . The amplitude of these changes is
larger in the top 1000 m, but anomalies below 1000 m span a large portion of the water column.
Correlation between transport components
During the first 3 years of observations (2004–2007), the
components of the MOC (FC, UMO, and Ekman) showed little co-variability,
leading to the conclusion that components contribute their variability
independently to the MOC . More
recently, sporadic periods of co-variability were identified between
currents at the western boundary: the FC and Antilles Current from 2009 to 2011, the end of the record at the time of
publication
. From the 10-year record
(2004–2014), we now see correlations emerging between some of the
contributing terms, which has implications for how we understand the
large-scale circulation at 26∘ N. Considering the 3-month-filtered UMO and FC transport time series (black lines in Fig. 2),
anomalies of opposite sign appear to coincide in late 2008, late 2010,
and again in late 2012. During the 2010/11 winter, for example, the
northward flowing FC weakened by several Sverdrups. At the same time,
the southward flowing UMO weakened by nearly 10 Sv. We investigate these
apparent compensations (between the UMO and FC, and also between LNADW
and wind-driven Ekman flow) more rigorously in the following section.
UMO and FC transports: Horizontal
circulation
The FC carries most of the waters of the Gulf Stream across 26∘ N.
The origins of this water come from the Yucatan Channel and Old Bahama
Channel, across complex topography west of the Bahamas
. At similar latitudes, the flow through the Yucatan
Channel has been found to compensate flow around Cuba ,
while variations in the FC have, at times, shown compensation east of the
Bahamas in bottom pressure variations and in top
1000 m velocities .
Transport anomaly time series (left column) for the -FC (blue) and
UMO (magenta). Zero anomaly is marked with the dashed line. Scatter plots of
the same (right column) with correlation coefficients (r) noted. In
(a, b), the deseasonalized, detrended time series are used. In
(c, d), the time series are deseasonalized and low-pass filtered,
but not detrended. The orthogonal regression line is overlaid on the scatter
plots. Grey bars highlight periods noted in the text.
Here we consider compensation between FC and UMO, the transbasin transport
east of the Bahamas. This compensation can be clearly seen by plotting their
detrended anomaly time series (Fig. 5a). Certain events stand out,
demonstrating almost perfect correspondence between the two time series, with
examples including February–May 2007, September 2008–June 2009, August
2010–January 2011 and August 2012–March 2013 (highlighted in the figure).
Notably, these episodes of correlation are absent in the first
3 years. The overall correlation between the two records is
r=-0.49, significant at the 95 % level.
Fluctuations in UMO compensate fluctuations in the FC by similar magnitudes
(slope=-0.92, Fig. 5b). When the northward FC transport increases
along the western boundary, the southward UMO transport east of the Bahamas
intensifies by the same amount. This means that excess northward flow in the
boundary current is returned horizontally within the upper mid-ocean
circulation rather than by deeper layers in the interior, which would have
involved changes in the MOC. The region east of the Bahamas is known to be
rich with eddies, which may influence the transbasin transports
,
and due to the timescale of observed compensation, we suspect that eddies are
involved.
Using the low-pass filtered time series, this high degree of compensation is
absent (Fig. 5c and d). Instead, strong interannual variability in the UMO
remains (Fig. 5c). By comparison, the low-pass filtered FC shows little
interannual variability, consistent with previous work that indicated that
the interannual and longer period FC variability is of much smaller amplitude
than the sub-annual variability (e.g.
). While the two time series are
not significantly correlated (r=0.24), both show a reduction from the first
5-year period (April 2004–March 2009) to the latter 5-year period
(April 2009–March 2014; see also Table 1), with the FC reducing by
0.7 Sv and the UMO by 1.9 Sv. Unlike the compensation at
higher frequencies, these changes are both of the same sign (note that the
negative of UMO is plotted in Fig. 5a), compounding the effect on the MOC
(net reduction of 2.9 Sv).
LNADW and Ekman transports: deep wind-driven
response
Another – and perhaps more remarkable – correlation that emerges from this
analysis is between the deepest limb of the southward mid-ocean
transports (LNADW) and the surface meridional Ekman transport (Fig. 6a).
Using the detrended anomaly time series, the typically southward
LNADW transport can be seen to reduce or even temporarily reverse to
northward during strong Ekman transport reversals. (See, e.g., events in
December 2009–April 2010, November 2010–January 2011, February–March 2013.) The correlation is
statistically significant (r=-0.58), and can be seen to occur
throughout much of the record rather than just during the extreme
events. In the low-pass filtered data (Fig. 6c and d), the correlation is
stronger than for the UMO and FC. However, due to the low number of
degrees of freedom for the low-pass filtered time series
(ndof=5), it is not significant.
As Fig. 5 but for the Ekman and -LNADW transport anomaly time
series.
As with the FC and UMO compensation, magnitudes of fluctuations between Ekman
and LNADW match (slope=-0.84, with Ekman anomalies of 1 Sv
corresponding to a 0.84 Sv change in the LNADW). Unlike the FC
and UMO, however, the correlation between the LNADW and Ekman at higher
frequencies projects onto the MOC rather than cancelling. This is consistent
with expectations that the high-frequency, wind-driven variability of Ekman
transport results in an overturning, albeit a shallow one (where the depth
of overturning is at the base of the northward Ekman transport) between the
surface Ekman transport and return flow below
. Variations in Ekman or FC project
directly on the mid-ocean transport (through the Text term in 2),
and bottom pressure records at the western boundary also covary with Ekman
anomalies . However, we will show that the
covariations between FC, Ekman, and mid-ocean transports are not limited to
the Text contribution, but are instead associated with density
changes at the western boundary.
To identify possible lags between the UMO and FC or Ekman and LNADW, we use
the 10-day filtered time series. For both correlations, between the
UMO and FC and between the Ekman and LNADW, the timescale of the response is
fast (Fig. 7). For the LNADW and Ekman correlation, a maximum correlation of
r=0.51 is found at 1-day lag with Ekman leading. This means that
the wind response occurs essentially instantaneously. For the UMO and FC
transports, a maximum correlation of r=0.46 is found for UMO leading by
0.5 days. This lead–lag relationship can also be seen by inspecting
close-zoomed plots of the time series during large anomalies (Fig. 7e). Due
to filtering applied to individual instrument data and transport time series,
such a short lag is not statistically meaningful.
10-day filtered transport anomaly time series of UMO and FC
(a, b) and LNADW and Ekman (c, d). The 10-year mean has
been removed, but the seasonal cycle and trends are retained. Time ranges
have been chosen to correspond to large anomalies in both time series, to
visualize possible lags. (e) Lag correlation between 10-day
filtered UMO and FC (magenta), and Ekman and LNADW (green); 95 %
significance is marked by the dashed lines, same colour.
Depth structure of
co-variability
The hypsometric compensation term (Text) is mostly
depth-independent, but has a vertical profile that scales with the width of
the basin as a function of depth. It is nearly uniform from the surface to
about 3500 m, and then decreases gradually to zero at the greatest
depths in the basin. If the mid-ocean region had no shear
(Tint=0) and no flow in the wedge (Twbw=0), the MO
transport would still be non-zero through this applied compensation, in order
to balance the northward FC and Ekman transport. In the absence of strong
variations in Tint, we would expect to see anti-correlation
between the MO transports (e.g. the UMO and LNADW) and the independently
estimated FC and Ekman transports. In this case, the MO transport
fluctuations would then have a depth structure approximately matching the
hypsometric profile. Instead, the MO transport-per-unit-depth profiles often
show deep maxima below 3500 m in anomaly plots (Fig. 3a).
The deep maximum in transport per unit depth implies that there are
considerable changes in the deep shear integrated across the width of the
basin, which are reflected in the Tint term. (There is little
transport in the Twbw term below 2000 m.) The Tint
term, in turn, is determined by the isopycnal displacements at the eastern
and western boundaries, according to a basin-wide thermal wind balance. In
the following, we will use isopycnal displacements at the two basin
boundaries to investigate the vertical structure of the variability noted
above. In the simplest case of a two-layer fluid, a tilted interface marked
by the displacement of the interface at the boundaries will have an
associated geostrophic shear between the layers. The strength of the shear
increases with the tilt of the interface so that if the interface were at
a constant depth at the eastern boundary, the shear would be controlled by
displacements at the western boundary.
While the stratification at 26∘ N is more continuous, we use
displacements at the two basin boundaries to investigate the magnitude of the
shear variability and the role of east and west in producing these shear
anomalies. Using the time-mean density gradient profile, given by
N‾2(z)=-(g/ρ)∂ρ¯/∂z, the contribution
of heave to transport per unit depth can be estimated from Eq. (2) as
T̃int(z)=1f∫zzrefzζeNe‾2-ζwNw‾2dz′,
where ∼ distinguishes this portion of Tint from the more
complete calculation in Eq. (2). Here, ζe(z,t) and
ζw(z,t) are the isopycnal displacements at the east and
west, respectively. From Eq. (3), we might expect some correspondence between
isopycnal displacements and transports, though transports are the
vertically integrated transport per unit depth.
Comparing the UMO transport with both eastern and western boundary isopycnal
displacement time series, we find strong correlations (Fig. 8a). In the west,
displacements between 300 and 1200 m are significantly correlated
with the UMO time series, with a peak at 820 m. This is consistent
with physical expectations, and prior results indicated a role for the
displacement of the main thermocline in controlling the UMO transport
. The
correlation in the east is of similar absolute amplitude but is spread over
a broader depth range (200 to ∼1800m deep), consistent with
the findings of that the eastern boundary
density variations were coherent down to 1400 m. No significant
correlations were found between any of the transport components considered
here (UMO, FC, LNADW, or Ekman) and the isopycnal displacements at the
Mid-Atlantic Ridge (not shown).
Correlation between isopycnal displacements at each depth and
transport time series, where time series are deseasonalized and detrended.
(a) Correlation between UMO (magenta) and isopycnal displacements at
the west (solid) and east (dashed), and between FC (blue) and isopycnal
displacements at the west (solid) and east (dashed). Significant correlations
are indicated by the thicker line. (b) The same, but for LNADW
(purple) and Ekman (green).
The FC is also highly correlated with the western boundary thermocline
displacement (Fig. 8a). The sign of the correlation has flipped, consistent
with the anti-correlation noted between the FC and UMO. This relationship is
statistically significant, even though the isotherms covarying with the FC
are 150 km away from the FC, east of the Bahamas. As might be
expected, there is no statistically significant relationship between the
thermocline displacements on the eastern side of the basin and the FC
transport. Comparing the time series of western boundary isopycnal
displacements at 820 m with the UMO transport (Fig. 9a), we find
significant correlation where a 10 m downward displacement of the
thermocline corresponds to a 1 Sv increase in the UMO transport
(Fig. 9b). Given the one-to-one relationship between the FC and UMO
(Fig. 5b), this means that a 10 m thermocline displacement is also
associated with a 1 Sv change in the FC.
Time series of deseasonalized, detrended transports and isopycnal
depths. (a) UMO transport (magenta) and the depth of the density
surface with mean position at 820 m (highest correlation with UMO) at
the west. (c) LNADW transport (purple) and the depth of the density
surface with highest correlation (at 3140 m). Scatter plots are shown
in (b, d), where the least-squares linear regression is overlaid.
The correlation between LNADW transport and isopycnal displacements is
significant at depth (1500 m–bottom) on the western boundary only
(Fig. 8b). The negative correlation can be readily understood in that
a depression of deep isopycnals on the western side of the basin causes
larger northward shear below 2000 m. That is to say, a depression of
deep isopycnals on the western side of the basin will result in a slowing of
the southward LNADW (a northward anomaly) relative to the UNADW layer above
it. One remarkable feature of this correlation is that it extends vertically
over several thousand metres of water (spanning several moored instruments),
so that when water at 3000 m moves upwards, a large segment of the
water column above and below is also moving upwards (though with differing
magnitudes). In this case, a 42 m downward displacement of the
isopycnal at 3140 m results in a 1 Sv reduction in the LNADW
transport (Fig. 9d).
Isopycnal displacements in a reduced region (2700–3300 m) at the
western boundary are also significantly correlated with the surface Ekman
transport (Fig. 8b). Positive (northward) Ekman transport anomalies are
associated with upward displacements of the deep isopycnals. This means that
when the winds blow along 26∘ N, the deep ocean responds by
heaving upwards or downwards across hundreds of metres, with the end result
that these isopycnal displacements at the western boundary change the
basin-wide tilt and thus the vertical shear in meridional transports. Just
40 km offshore (25 km further offshore than the western
boundary) at the WB3 mooring, isopycnal displacements are still significantly
correlated with LNADW transports, albeit more weakly (|r|≤0.5, not
shown). Ekman transports are no longer correlated with displacements. At the
WB5 mooring, 500 km offshore, there is no relationship between
isopycnal displacements and basin-wide transport. The strong correlation
between isopycnal displacements nearshore and meridional Ekman transport, and
the absence of correlation for offshore displacements may indicate that the
deep compensation is concentrated at the western boundary, or that other
variability in isopycnal displacements masks the signal offshore.
Timescales of
compensation/co-variability
One of the key results presented here is that the UMO and FC transports
often compensate each other – i.e. their signs differ but anomalies match – resulting in greatly reduced
impact of their individual fluctuations on the total MOC variability.
However, this compensation is dependent on timescales. At
low-frequencies, the compensation does not dominate (Fig. 5c), and the large
interannual variability and trend in the UMO transport has a strong
projection onto the interannual variability and trend of the MOC.
To investigate the co-variability for different timescales, we evaluate the
coherence calculated using a multitaper spectrum following
(Fig. 10). The FC and UMO are significantly
coherent and out-of-phase (i.e. anti-correlated) at periods less than
1 year. For periods longer than about a year they are no longer
coherent. For Ekman and LNADW transports, they are coherent at periods
shorter than 900 days (except near 120 days) and also
(nearly) out-of-phase. By contrast, there is little coherence between the
Ekman and FC time series (Fig. 10, grey).
Coherence between MOC components: UMO and FC (magenta), LNADW and
Ekman (green) and FC and Ekman (grey), where time series are the original
10-day filtered (seasonal variations retained). The top panel shows
coherence, where significance is delimited by the black horizontal line. The
lower panel shows the phase relationship at each period in degrees.
These results at first appear to contradict , who
noted from the first 3 years of observations (2004–2007) that there
was no compensation between the FC and UMO. Using a cross-wavelet transform
, we can examine the temporal shifts in
co-variability between MOC components (Fig. 11). Since the beginning of
2008, the FC and UMO covary at annual to sub-annual timescales (consistent
with the coherence calculation). Prior to 2008 there is no significant
covariation between them (Fig. 11a), consistent with the findings of
. The cross-wavelet transform of LNADW and Ekman
shows that the periods of higher co-variability occur in the latter half of
the 10-year record (Fig. 11b). There also exists some power at annual
timescales and higher frequencies (e.g. during the wind-reversal events in
2005, 2009/10, and 2013).
(a) Cross-wavelet transform between the FC and UMO shows
high power (red) with a fixed-phase relationship since 2007 at periods
between 60 and 400 days. (b) Cross-wavelet transform between
Ekman and LNADW shows high power primarily at annual periods during the
2009–2010 events, as well as sub-annual from 50–150 days in 2005,
2011, and 2013. (Arrows pointing to the left indicate out-of-phase
relationship or anti-correlation, and are only shown when the relationship
is significant. Deviations from 180 ∘ indicate a lag or phase
shift.)
The compensation between the FC and UMO is non-stationary. A windowed
correlation between the FC and UMO prior to 2007 shows no significant
correlation, either between the full time series or the high-pass filtered
time series (not shown). From , the FC and
Antilles Current (AC) were correlated between 2009 and 2011. The AC is a
narrow, northward boundary current east of the Bahamas. Its transport is
measured by direct current metre measurements in the RAPID array, and is
included in the UMO transport. The mean transport of the AC is relatively
small (3.6 Sv, compared to 31 Sv in the FC), so that the UMO transport is
still net southward. However, since the AC is part of the UMO transport, a
correlation between the FC and AC should weaken an anti-correlation between
the FC and UMO. The isopycnal displacements associated with UMO and FC
variability (Fig. 8) also influence transport in the AC, as long as the AC is
geostrophic. The mooring used to estimate isopycnal displacements is at the
western edge of the UMO transport, but is east of the core of the northward
AC velocities. This means that a downward displacement of isopycnals is
associated with an increase in the tilt of the thermocline between Africa and
the mooring, but will have an opposing effect on geostrophic transports west
of the mooring, in the Antilles Current.
identified stronger eddy activity east of the Bahamas during the period of FC
and AC correlation, which is consistent with a eddy activity playing a role
in the anti-correlation between the FC and UMO on subannual timescales.
Trends in isopycnal displacements and MO
transports
The MOC at 26∘ N has been decreasing in strength, as described in
Smeed et al. (2014) and this trend continues through early 2014. This low-frequency change is mainly associated with changes in the UMO
transport (variations at timescales longer than 1 year) and to
a lesser degree with a weak reduction in the FC transport (Table 1). From the
low-pass filtered transport-per-unit depth profiles (Fig. 3b), we noted that
changes are present both above and below the thermocline. Here we investigate
how the trends in transport are captured by trends in isopycnal
displacements.
In the top 1000 m, the isopycnal slope is large, with lighter water
in the west. The slope decreases to 0 at about 1000 m, reversing sign
between 1000–1500 m, and is typically small below (Fig. 12a). At
the western boundary, isopycnal displacements below 1000 m are moving
downwards (trend at 3140 m is -6.5±3.5myr-1,
increasing to about -13±0.4myr-1 at 4500 m),
while above 500 m they are moving upwards (Fig. 12b). At the
eastern boundary, trends are near zero except below 4000 m where they
are downward. As a result, the east–west slope of isopycnals below
1000 m is decreasing with time. While the largest displacements are
seen at depth, stratification is also weaker at depth. From Eq. (3), we see
that the effect of isopycnal displacements on transports is modulated by the
background stratification. Computing instead the trends in isopycnal
displacements scaled by stratification (Fig. 12c), we now see that at the
western boundary, the effect of isopycnal displacements below the thermocline
is nearly constant, with little effect on the eastern boundary. Scaling by
stratification emphasizes
the importance of relatively small trends in
displacements in the top 1000 m to transport anomalies.
Trend in isopycnal displacements and transports. (a) Mean
height difference between isopycnals in the east and west. (b) Trend
in isopycnal displacements at the western and eastern boundaries, where
a negative (positive) trend indicates downward (upward) movement of
isopycnals. The solid (dashed) line is for displacements at the west (east).
(c) Trend in displacements scaled by buoyancy frequency with
N0=1×10-3s-1. (d) Trend in
T̃int from Eq. (2). The red line is the trend in
MO(z)-Text. (e) Trends in
transport per unit depth of mid-ocean (MO(z)) in magenta,
Text in blue. All time series were twice monthly. Confidence
intervals on the trends are shaded. The dotted line shows a zero trend.
The trend in T̃int from Eq. (3) is shown in Fig. 12d,
with a persistent reducing trend in the interior transports associated
with heave at the eastern and western boundaries. The magnitude of the
trend increases from 0 at the bottom, since it is integrated upwards
from the bottom. While the thermocline displacement shows little-to-no
trend (Fig. 12b, the trend on the western boundary at 830 m is
-0.7±0.9myr-1), the interior transport has a large
negative trend. The large amplitude at the depth of the thermocline results
from the accumulation of persistent negative anomalies between the
bottom and the thermocline. Above 1000 m, the trend in T̃int(z) is
relatively constant and negative.
A southward intensification of Tint, if unbalanced by
other components, would result in an overall intensification of the southward
flow across the section over the 10-year period. To maintain overall mass
balance across the section, an opposing trend is required in Text
(Fig. 12e). Trends in the overall mid-ocean transport (MO(z))
include those from T̃int as well as anomalies in the wedge
and compensation, due to stratification changes and diabatic changes, and both at the boundaries and Mid-Atlantic Ridge. The overall amplitude
of the trend in MO(z) is weaker than in T̃int(z), but
with a greater shear near the surface. The wedge in particular
contributes to the shear in the top 1000 m in MO(z) relative to
T̃int (not shown).
The trend in MO(z) is negative above ∼1700m and
positive below. This amounts to a rebalancing of the absolute reference
level for Tint to account for the changing structure of the
interior ocean shear profile. Summing the Tint and Text
components results in a northward anomaly below 1700 m (where
Text exceeds anomalies in Tint). The vertical structure of
these trends is consistent with the intensification of the baroclinic circulation: the
LNADW is reducing in intensity (less southward than before) while the UMO is
increasing in southward flow. In addition, the depth of the change in the trend (1700 m) is
in the middle of the UNADW layer (1100–3000 m) offering an explanation
for why no long-term trend is apparent in the UNADW transport.
Discussion
Here we have identified significant compensation, dependent on timescale,
between components of the MOC: for example, on sub-annual timescales when
the FC is stronger northward in the western boundary, the UMO compensates
with stronger southward flow between the Bahamas and Canary islands. While
these components are largely independent, they are weakly coupled due to the
construction of the MO(z) transports. The Text term
contributes variability to both the UMO and LNADW transport, and would tend
to cause an opposite sign anomaly in both the UMO and LNADW in response to
anomalies in the FC and Ekman transports.
However, by construction the Text term is nearly barotropic, and
what limited vertical structure it has arises from the vertical profile
of section width at 26∘ N. This means that Text
distributes compensation for the Ekman or FC anomalies across the full
water column, with the strongest compensation in the MO(z) term
occurring in the top 3500 m and reducing to zero at the bottom. Only
a fraction of FC anomalies (order one-fifth) would be projected by Text
onto UMO anomalies, and less than half the magnitude of Ekman anomalies
would be contained in the LNADW portion of the Text. Instead, the
observed transport anomalies in UMO and LNADW are indistinguishable in
magnitude to the anomalies in FC and Ekman transports. Furthermore,
there is no correlation between the Ekman and UNADW transport, which
requires a shear between LNADW and UNADW to remove any projected
compensation to Ekman. We can see this baroclinic response in the
isopycnal displacements measured at the western boundary.
The compensation between the UMO and FC has the appearance of
a time-varying horizontal or gyre circulation, but is limited to
sub-annual timescales similar to those of eddies identified in
. In that paper, on some
occasions, when a large cyclone (anticyclone) impacted the Bahamas (and
Twbw transports), the FC would respond by weakening
(strengthening). Sea surface height anomalies originating east of the
Bahamas coincided with same sign anomalies along the eastern side of the
Florida Straits, resulting in transport anomalies of opposite sign to
the west and east of the height anomalies. The new result
here is that we use the full UMO rather than just the Antilles Current
as in the 2013 paper, and the pattern of anti-correlation has persisted since 2007.
Because the co-variability of the FC and UMO transports is accompanied
by roughly equal magnitude anomalies, there is a reduced projection of
either component on the variability on the overall MOC. This is also
consistent with the findings of , who used the RAPID data and sea surface height anomalies to show that Rossby
waves/eddies with a sub-annual timescale do contribute to dynamic height
anomalies at the boundary but do not affect the MOC.
The relationship between surface winds and deep isopycnal displacements
is harder to understand, and we presently do not have a dynamical
explanation for this behaviour. Nevertheless, the observations are quite
clear: when there is anomalous southward Ekman transport (resulting from
westerly winds), isopycnals at ∼3000m on the western
boundary plunge downwards. From theoretical and numerical model results
, it is
expected that Ekman transport fluctuations on sub-annual timescales
should result in a quasi-instantaneous, and nearly barotropic,
compensation in the deep ocean.
One might expect that with anomalous
southward Ekman transport across the basin, there should be a northward
compensation across a broad range of depths, which could also perhaps be
spatially distributed. showed in a numerical
model that Ekman anomalies project strongly onto the barotropic mode
promoting wave interactions with deep bathymetry. It is possible that
the baroclinic response we observe at 26∘ N is generated by
these barotropic waves interacting with the bathymetry at the western
boundary. This is an area for future research – likely requiring
numerical modelling to isolate mechanisms and processes.
The observed high-degree of variability on sub-annual timescales of the
UMO and LNADW transports may contribute to the apparent absence of
meridional coherence between observations at 26∘ N and higher
latitudes. found large coherence on
short timescales at mid-latitudes in the North Atlantic, but between
41 and 26∘ N, transport anomalies were out-of-phase
. The
large fluctuations in LNADW transport are remarkably well captured by ocean
bottom pressure from the GRACE satellites .
However, the spatial footprint of detrended monthly anomalies is centred in
the western part of the subtropical North Atlantic, suggesting limited
meridional coherence. Using satellite altimetry,
found that sea surface height anomalies capture the interannual variability
and trend of the UMO transport, with a spatial footprint extending over
a larger area.
Conclusions
The record of basin-wide MOC transport variability at 26∘ N in the
Atlantic is now 10-years long. It continues to deliver new insights into the
origins of the changing large-scale circulation at 26∘ N. In this
paper, we have provided a brief overview of the latest 18 months of
observations. In particular, with the most recent 18 months of
observations, the April 2012 to March 2013 year has the second
weakest MOC, behind the 2009/10 year. Unlike the
2009/10 year, the reduction in 2012/13 is associated with an enhanced
southward UMO and decreased LNADW, with no contribution from anomalous Ekman
transport. There were, however, several wind-reversal events in the latest
18 months, at the end of 2012 and again in March 2013. These events
were shorter than the 2009/10 and 2010/11 double dip, though they also
resulted in concurrent reductions in the southward LNADW transport.
The main result of this paper has been detailing newly identified
compensations between MOC components (UMO and FC, and LNADW and Ekman).
Using the 10-year record, we now find that on shorter timescales (periods
shorter than 1 year), much of the variability of the UMO is
compensated by the FC transport variability, particularly in the most recent
7 years. On similar timescales (periods shorter than 2 years),
the wind-driven variability in the top 100 m (surface meridional
Ekman transport) is nearly instantaneously balanced by deep flow in the
opposite direction. However, rather than being a simple barotropic response
to the winds, the imprint of the winds is baroclinic, with the strongest
signature in isopycnal displacements being found just east of the Bahamas
at 3000 m depth.
There is a key difference between these two compensating transport pairs.
Between the FC and UMO, compensated high-frequency transport anomalies
results in a horizontal circulation anomaly: that is, northward flow in the
FC (in the top 700 m) is accompanied by southward flow in the top
1100 m east of the Bahamas. As a consequence, transport anomalies
have little influence on the MOC and are unlikely to have a strong heat
transport anomaly. By contrast, the southward Ekman anomalies accompanied by
northward LNADW anomalies directly projects onto the MOC. During the
anomalous periods of exceptionally strong (weak) Ekman transport, the MOC is
similarly strong (weak) and we expect the meridional heat transport to vary
with the MOC .
Finally, investigating longer-term variations of the MOC, we can localize
the origin of the intensifying trend in the MOC to isopycnal displacements on
the western boundary. Observed transport fluctuations on interannual and
longer timescales present a different story. From the 10-year record, the
mid-ocean transports (rather than FC and Ekman) are primarily responsible for
the low-frequency variability and trend of the overturning. Furthermore, the
trend in transport variability is associated with the persistent deepening of
isopycnals below the thermocline at the western boundary. These displacements
are greatest in the abyss (130±40m over 10 years at
4500 m) compared to about 60±30m at mid-depths around
2000 m, though their impact on transports must be scaled by
stratification. While we do not investigate here whether the longer-term
isopycnal deepening is associated with water mass changes or wind-forcing,
a coincident shift towards warmer and fresher waters below 3000 m hints at the possibility of larger-scale
persistent changes to the Atlantic circulation.
Acknowledgements
Data from the RAPID Climate Change
(RAPID)/Meridional overturning circulation and heat flux array (MOCHA), Western
Boundary Time Series (WBTS) projects are funded by the Natural
Environment Research Council (NERC), National Science Foundation (NSF,
OCE1332978) and National Oceanic and Atmospheric
Administration (NOAA), the Climate Program Office – Climate Observation Division. Data are freely available from
www.rapid.ac.uk.
Florida Current transports are funded by the
NOAA and are available
from
www.aoml.noaa.gov/phod/floridacurrent.
Wavelet code provided by A. Grinsted, J. Moore and S. Jevrejeva. Special
thanks to the captains, crews, and technicians, who have been invaluable in the
measurement of the MOC at 26∘ N over the past 10 years.
Edited by: M. Hoppema
ReferencesAtkinson, C. P., Bryden, H. L., Cunningham, S. A., and King, B. A.: Atlantic
transport variability at 25∘ N in six hydrographic sections, Ocean
Sci., 8, 497–523,
doi:10.5194/os-8-497-2012, 2012.Bryden, H. L., Longworth, H. R., and Cunningham, S. A.: Slowing of the Atlantic
meridional overturning circulation at 25∘ N, Nature, 438,
655–657, 2005.Bryden, H. L., Mujahid, A., Cunningham, S. A., and Kanzow, T.: Adjustment of the basin-scale
circulation at 26∘ N to variations in Gulf Stream, deep western
boundary current and Ekman transports as observed by the Rapid array, Ocean
Sci., 5, 421–433,
doi:10.5194/os-5-421-2009,
2009.Bryden, H. L., King, B. A., McCarthy, G. D., and McDonagh, E. L.: Impact of a 30 %
reduction in Atlantic meridional overturning during 2009–2010, Ocean Sci.,
10, 683–691,
doi:10.5194/os-10-683-2014,
2014.Cabanes, C., Lee, T., and Fu, L.-L.: Mechanisms of interannual variations of
the Meridional Overturning Circulation of the North Atlantic, J. Phys.
Ocean., 38, 467–480, 10.1175/2007JPO3726.1, 2008.Chidichimo, M. P., Kanzow, T., Cunningham, S. A., Johns, W. E., and Marotzke,
J.: The contribution of eastern-boundary density variations to the Atlantic
meridional overturning circulation at 26.5∘ N, Ocean Sci., 6,
475–490,
doi:10.5194/os-6-475-2010, 2010.Clément, L., Frajka-Williams, E., Szuts, Z. B., and Cunningham, S. A.: The vertical
structure of eddies and Rossby waves and their effect on the Atlantic MOC at
26∘ N, J. Geophys. Res., 119, 6479–6498,
doi:10.1002/2014JC010146,
2014.Cunningham, S. A., Kanzow, T., Rayner, D., Baringer, M. O., Johns, W. E., Marotzke, J.,
Longworth, H. R., Grant, E. M., Hirschi, J. J.-M., Beal, L. M.,
Meinen, C. S., and Bryden, H. L.: Temporal variability of the Atlantic
meridional overturning circulation at 26.5∘ N, Science, 317,
935–938, 2007.Cunningham, S. A., Roberts, C., Frajka-Williams, E., Johns, W. E., Hobbs, W., Palmer, M. D.,
Rayner, D., Smeed, D. A., and McCarthy, G. D.: Atlantic MOC slowdown cooled
the subtropical ocean, Geophys. Res. Lett., 40, 6202–6207,
doi:10.1002/2013GL058464,
2014.
Desaubies, Y. and Gregg, M. C.: Reversible and irreversible finestructure, J.
Phys. Ocean., 11, 541–556, 1981.
DiNezio, P. N., Gramer, L. J., Johns, W. E., Meinen, C. S., and Baringer, M. O.: Observed
interannual variability of the Florida Current: Wind forcing and the North
Atlantic Oscillation, J. Phys. Oceanogr., 39, 721–736, 2009.Duchez, A., Frajka-Williams, E., Castro, N., Hirschi, J. J.-M., and Coward, A.: Seasonal to
interannual variability in density around the Canary Islands and their
influence on the AMOC at 26.5∘ N, J. Geophys. Res., 119,
1843–1860,
doi:10.1002/2013JC009416,
2014.
Elipot, S., Hughes, C., Olhede, S., and Toole, J.: Coherence of western boundary pressure at
the RAPID WAVE array: Boundary wave adjustments or deep western boundary
current advection?, J. Phys. Oceanogr., 43, 744–765, 2013.Elipot, S., Frajka-Williams, E., Hughes, C., and Willis, J.: The observed North Atlantic MOC,
its meridional coherence and ocean bottom pressure, J. Phys. Oceanogr., 44,
517–537,
doi:10.1175/JPO-D-13-026.1,
2014.
Emery, W. J. and Thomson, R. E.: Data Analysis Methods in Physical Oceanography, Elsevier,
Amsterdam, the Netherlands, 2nd edn., 2004.Frajka-Williams, E.: Estimating the Atlantic MOC at 26∘ N using satellite altimetry
and cable measurements, Geophys. Res. Lett., 42, 3458–3464,
doi:10.1002/2015GL063220,
2015.Frajka-Williams, E., Johns, W. E., Meinen, C. S., Beal, L. M., and Cunningham, S. A.: Eddy
impacts on the Florida Current, Geophys. Res. Lett., 40, 349–353,
doi:10.1002/grl.50115, 2013.Grinsted, A., Moore, J. C., and Jevrejeva, S.: Application of the cross wavelet
transform and wavelet coherence to geophysical time series, Nonlin. Processes
Geophys., 11, 561–566,
doi:10.5194/npg-11-561-2004,
2004.
Jayne, S. R. and Marotzke, J.: The dynamics of ocean heat transport variability, Rev. Geophys., 39, 385–411, 2001.Johns, W. E., Baringer, M. O., Beal, L. M., Cunningham, S. A., Kanzow, T., Bryden, H. L.,
Hirschi, J. J.-M., Marotzke, J., Meinen, C. S., Shaw, B., and Curry, R.:
Continuous, array-based estimates of Atlantic Ocean heat transport at
26.5∘ N, J. Climate, 24, 2429–2449, 2011.Kanzow, T., Cunningham, S. A., Rayner, D., Hirschi, J. J.-M., Johns, W. E., Baringer, M. O.,
Bryden, H. L., Beal, L. M., Meinen, C. S., and Marotzke, J.: Observed flow
compensation associated with the MOC at 26.5∘ N in the Atlantic,
Science, 317, 938–941, 2007.
Kanzow, T., Johnson, H. L., Marshall, D. P., Cunningham, S. A., Hirschi, J. J.-M., Mujahid, A.,
Bryden, H. L., and Johns, W. E.: Basinwide Integrated Volume Transports in an
Eddy-Filled Ocean, J. Phys. Oceanogr., 39, 3091–3110, 2009.Kanzow, T., Cunningham, S. A., Johns, W. E., Hirschi, J. J.-M., Marotzke, J., Baringer, M. O.,
Meinen, C. S., Chidichimo, M. P., Atkinson, C., Beal, L. M., Bryden, H. L.,
and Collins, J.: Seasonal variability of the Atlantic meridional overturning
circulation at 26.5∘ N, J. Climate, 23, 5678–5698,
doi:10.1175/2010JCLI3389.1,
2010.
Killworth, P. D.: A simple linear model of the depth dependence of the wind-driven
variability of the Meridional Overturning Circulation, J. Phys. Oceanogr.,
38, 492–502, 2008.Landerer, F. W., Wiese, D. N., Bentel, K., Boening, C., and Watkins, M. M.: North
Atlantic meridional overturning circulation variations from GRACE ocean
bottom pressure anomalies, Geophys. Res. Lett.,42, 8114–8121,
doi:10.1002/2015GL065730,
2015.Lin, Y., Greatbatch, R. J., and Sheng, J.: A model study of the vertically integrated
transport variability through the Yucatan Channel: Role of Loop Current
evolution and flow compensation around Cuba, J. Geophys. Res., 114, C08003,
doi:10.1029/2008JC005199,
2009.Longworth, H. R., Bryden, H. L., and Baringer, M. O.: Historical variability in
Atlantic meridional baroclinic transport at 26.5∘ N from boundary
dynamic height observations, Deep-Sea Res. Pt. II, 58, 1754–1767, 2011.McCarthy, G., Frajka-Williams, E., Johns, W. E., Baringer, M. O., Meinen, C. S.,
Bryden, H. L., Rayner, D., Duchez, A., Roberts, C. D., and Cunningham, S. A.:
Observed interannual variability of the Atlantic MOC at 26.5∘ N,
Geophys. Res. Lett., 39, L19609,
doi:10.1029/2012GL052933,
2012.McCarthy, G. D., Smeed, D. A., Johns, W. E., Frajka-Williams, E., Moat, B. I.,
Rayner, D., Baringer, M. O., Meinen, C. S., and Bryden, H. L.: Measuring the
Atlantic meridional overturning circulation at 26∘ N, Prog.
Oceanogr., 130, 91–111,
doi:10.1016/j.pocean.2014.10.006,
2015.
Meinen, C. S., Baringer, M. O., and Garcia, R. F.: Florida Current transport variability:
An analysis of annual and longer-period signals, Deep-Sea Res. Pt. I, 57,
835–846, 2010.Mielke, C., Frajka-Williams, E., and Baehr, J.: Observed and simulated variability
of the AMOC at 26∘ N and 41∘ N, Geophys. Res. Lett.,
40, 1159–1164,
doi:10.1002/grl.50233, 2013.
Percival, D. B. and Walden, A. T.: Spectral Analysis for Physical Applications,
Cambridge University Press, Cambridge, UK, 1998.Pillar, H., Heimbach, P., Johnson, H., and Marshall, D.: Dynamical attribution
of recent variability in Atlantic overturning, J. Climate,
10.1175/JCLI-D-15-0727.1, 2016.Polo, I., Robson, J., Sutton, R., and Bamaseda, M.: The importance of wind
and buoyancy forcing of the boundary density variations and the geostrophic
component of the AMOC at 26∘ N, J. Phys. Oceanogr., 44, 2387–2408, 2014.Rayner, D., Hirschi, J. J.-M., Kanzow, T., Johns, W. E., Wright, P. G., Frajka-Williams, E.,
Bryden, H. L., Meinen, C. S., Baringer, M. O., Marotzke, J., Beal, L. M., and
Cunningham, S. A.: Monitoring the Atlantic meridional overturning
circulation, Deep-Sea Res. Pt. II, 58, 1744–1753,
doi:10.1016/j.dsr2.2010.10.056,
2011.Roberts, C. D., Waters, J., Peterson, K. A., Palmer, M., McCarthy, G. D.,
Frajka-Williams, E., Haines, K., Lea, D. J., Martin, M. J., Storkey, D.,
Blockley, E. W., and Zuo, H.: Atmosphere drives observed interannual
variability of the Atlantic meridional overturning circulation at
26.5∘ N, Geophys. Res. Lett., 40, 5164–5170,
doi:10.1002/grl.50930, 2013.Roberts, C. D., Jackson, L., and McNeall, D.: Is the 2004–2012 reduction of
the Atlantic meridional overturning circulation significant?, Geophys. Res.
Lett., 41, 3204–3210,
doi:10.1002/2014GL059473,
2014.
Robson, J., Hodson, D., Hawkins, E., and Sutton, R.: Atlantic overturning in decline?,
Nat. Geosci., 7, 2–3, doi:10.1038/ngeo2050,
2014.Rousset, C. and Beal, L. M.: Closing the transport budget of the Florida Straits, Geophys. Res. Lett., 41, 2460–2466,
doi:10.1002/2014GL059498, 2014.Smeed, D. A., McCarthy, G. D., Cunningham, S. A., Frajka-Williams, E., Rayner,
D., Johns, W. E., Meinen, C. S., Baringer, M. O., Moat, B. I., Duchez, A., and
Bryden, H. L.: Observed decline of the Atlantic meridional overturning
circulation 2004–2012, Ocean Sci., 10, 29–38,
doi:10.5194/os-10-29-2014,
2014.Smeed, D. A., McCarthy, G., Rayner, D., Moat, B. I., Johns, W. E.,
Baringer, M. O., and Meinen, C. S.: Atlantic meridional overturning
circulation observed by the RAPID-MOCHA-WBTS (RAPID-Meridional Overturning
Circulation and Heatflux Array-Western Boundary Time Series) array at
26∘ N from 2004 to 2014, available at: https://www.bodc.ac.uk/data/published_data_library/catalogue/10.5285/1a774e53-7383-2e9a-e053-6c86abc0d8c7/, last access: 15 August 2015.Thomas, M. D. and Zhai, X.: Eddy-induced variability of the meridional overturning
circulation in a model of the North Atlantic, Geophys. Res. Lett., 40,
2742–2747,
doi:10.1002/grl.50532, 2013.Wunsch, C.: Mass and volume transport variability in an eddy-filled ocean, Nat. Geosci., 1, 165–168,
doi:10.1038/ngeo126, 2008.Wunsch, C. and Heimbach, P.: Two decades of the Atlantic meridional overturning
circulation: Anatomy, variations, extremes, prediction, and overcoming its
limitations, J. Climate, 26, 7167–7186,
doi:10.1175/JCLI-D-12-00478.1,
2013.Xu, X., Chassignet, E. P., Johns, W. E., Schmitz Jr., W. J., and Metzger, E. J.: Intraseasonal to interannual variability
of the Atlantic meridional overturning circulation from eddy-resolving
simulations and observations, J. Geophys. Res., 119, 5140–5159,
doi:10.1002/2014JC009994,
2014.Yang, J.: Local and remote wind stress forcing of the seasonal variability of the
Atlantic Meridional Overturning Circulation (AMOC) transport at
26.5∘ N, J. Geophys. Res., 120, 2488–2503,
doi:10.1002/2014JC010317,
2015.Yeager, S.: Topographic coupling of the Atlantic overturning and gyre circulations, J. Phys. Oceanogr., 45, 1258–1284,
doi:10.1175/JPO-D-14-0100.1, 2015.Zhao, J. and Johns, W.: Wind-forced interannual variability of the Atlantic
Meridional Overturning Circulation at 26.5∘ N, J. Geophys. Res.,
119, 2403–2419, 10.1002/2013JC009407, 2014.