Deep water formation, the subpolar gyre, and the meridional overturning circulation in the subpolar North Atlantic

Deep-Sea Research II 58 (2011) 1819–1832

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Deep water formation, the subpolar gyre, and the meridional overturning circulation in the subpolar North Atlantic a,b ¨ Monika Rhein a,n, Dagmar Kieke a, Sabine Huttl-Kabus , Achim Roessler a, Christian Mertens a, a b c ¨ , Igor Yashayaev d Robert Meissner , Birgit Klein , Claus W. Boning a

Institut f¨ ur Umweltphysik, Abt. Ozeanographie, Universit¨ at Bremen, Postfach 330440, 28344 Bremen, Germany Bundesamt f¨ ur Seeschifffahrt und Hydrographie, Bernhard-Nocht-Straße 78, 20359 Hamburg, Germany c Leibniz Institut f¨ ur Meereswissenschaften Ifm-Geomar, 24105 Kiel, Germany d Fisheries and Oceans Canada, Bedford Institute of Oceanography, Dartmouth, Nova Scotia, Canada b

a r t i c l e i n f o

abstract

Article history: Received 12 October 2010 Accepted 12 October 2010 Available online 31 January 2011

On interannual to decadal times scales, model simulations suggest a strong relationship between anomalies in the deep water formation rate, the strength of the subpolar gyre, and the meridional overturning circulation in the North Atlantic. Whether this is valid, can only be confirmed by continuous, long observational time series. Several measurement components are already in place, but crucial arrays to obtain time series of the meridional volume and heat transport in the subpolar North Atlantic are still missing. Here we summarize the recent developments of the deep water formation rates and the subpolar gyre transports. We discuss how existing observational components in the subpolar North Atlantic could be supplemented to provide long-term monitoring of the meridional heat and volume transport. Through a combined analysis of observations and model results the temporal and spatial scales that had to be covered with instruments are discussed, together with the key regions with the highest variability in the velocity and temperature fields. & 2011 Elsevier Ltd. All rights reserved.

Keywords: Overturning circulation Deep water formation Subpolar gyre transport

1. Introduction The subpolar North Atlantic is one of the climate relevant key regions of the ocean (Fig. 1). Here, the deep water masses of the cold limb of the Atlantic meridional overturning circulation (AMOC) are formed or significantly modified. An important element of this circulation is the Deep Western Boundary Current (DWBC), which provides a fast track for climate-induced anomalies to travel from the subpolar surface ocean into the tropical abyss (e.g. Fine et al., 2002). The upper limb of the AMOC consists of warm and salty water of tropical/subtropical origin, which is transported by the Gulf Stream into the mid-latitudes of the ¨ northern hemisphere (e.g. Kase and Krauss, 1996). After the separation from the continental slope the Gulf Stream – now dubbed North Atlantic Current (NAC) – branches out into several current bands, eddies and meanders, and crosses the Mid-Atlantic Ridge (MAR) into the eastern Atlantic. Some of this water flows into the Nordic Seas, but the major part turns westward, forming the anticlockwise subpolar gyre. The NAC subsequently releases heat to the atmosphere, providing a heat source for the relatively mild winters in Western Europe. The buoyancy loss of the NAC by

n

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the subsequent cooling along its anticlockwise pathway ultimately leads to the formation of deep and intermediate water in the Labrador Sea. This water transformation is one of the major connections between the warm upper branch and the cold lower branch of the AMOC. An important process for the intensity of this link is the energy exchange between atmosphere and surface ocean. In high-resolution ocean and climate models the intensity of deep water formation in the Labrador Sea and changes in the strength of the AMOC are closely linked on interannual to decadal time scales (e.g. Biastoch et al., 2008; Gregory et al., 2005; Stouffer et al., 2006). This relationship was first discussed by H¨akkinen ¨ (2001). Model studies by Boning et al. (2006) reveal a connection between deep water formation, AMOC, and the strength of the subpolar gyre, although the bulk of the variability of the gyre variation is caused by air-sea buoyancy flux and wind stress anomalies. Other climate models do not show any connection between these components (e.g. Fig. 5 in Landerer et al., 2007). To ascertain whether and how strongly related deep water formation, AMOC transport, and gyre strength are, long observational time series of the respective transports and formation rates are needed. At present, the observational basis is rather thin. Rhein et al. (2002) and Kieke et al. (2006, 2007) provided time series of Labrador Sea Water (LSW) formation for the period 1997–2003. Curry and McCartney (2001) provided a baroclinic

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transport index for the years from 1950 to 1997, representing the combined subpolar and subtropical gyre transport between Bermuda and the central Labrador Sea. This index was extended ¨ by Kieke et al. (2007) until 2003. Hakkinen and Rhines (2004) used time series of SSH anomalies to infer the weakening of the subpolar gyre in the 1990s, and using time series of current meters in the Labrador Sea linked it to the weakening of the LSW formation. Climate model scenarios predict a gradual slowdown of the AMOC in the 21st century as a consequence of global warming (Meehl et al., 2007). In this view and regarding the response it will have on climate and sea level rise at the west European coast (e.g. Levermann et al., 2005), it is highly desirable to expand the number of existing observations in the subpolar North Atlantic. Another motivation is how the changing circulation and water mass formation might affect the importance of the North Atlantic in storing anthropogenic carbon (Steinfeldt et al., 2009). In this study an update of the time series of the LSW formation rates and the Curry and McCartney (2001) transport index will be presented together with the recent observations of the water mass formation in the Labrador Sea. Up to now, measured time series of subpolar gyre transports outside the western boundary do not exist. In this study we present the first continuous time series of the subpolar gyre transport calculated from geostrophic moorings at the MAR for the years 2006–2008. Also lacking are time series of meridional volume and heat transports in the subpolar North Atlantic. Processes important for these transports are analyzed as they appear in the 1/121 ocean model FLAME. A particular focus will be on the latitude of 471N, and we will show how the observational basis at this latitude will improve in the next years. Finally, suggestions are discussed regarding how existing observational components could be expanded to continuously monitor the meridional volume and heat transport in the subpolar North Atlantic.

2. Database and applied methods 2.1. Deep water formation rates inferred from tracer inventories Rhein et al. (2002) and Kieke et al. (2006, 2007) calculated the chlorofluorocarbon (CFC) inventories of Labrador Sea Water, and inferred from biennial changes of these inventories LSW formation rates for the years 1997–2003. Calculations were done for two modes of LSW, the upper LSW (ULSW, bounded by sY ¼27.68– 27.74 kg/m3) and the denser mode, sometimes called ‘deep’ LSW (sY ¼ 27.74–27.80 kg/m3). The latter is the product of the intense convection activity of the early 1990s penetrating down to about 2000 m (e.g. Lazier et al., 2002) and was not renewed between 1997 and 2003 (Kieke et al., 2007). The lighter ULSW is the product of the shallow convection reaching down to about 1500 m, and the respective formation rates decreased substantially until 2003 (Kieke et al., 2007). This time series is extended here by analyzing changes in the CFC-12 inventory of LSW between 2003 and 2005. Rhein et al. (2002) and Kieke et al. (2006) demonstrated a linear correlation between high CFC concentrations and low salinities in the LSW layers. Due to a coarser spatial coverage of tracer data in 2005 compared to earlier years, tracer concentrations were partly reconstructed from such linear correlations between CFC-12 and salinity. This approach was tested successfully on the inventories from previous years, which had a much higher spatial coverage. The additional uncertainties of the calculated inventories associated with this method were of the order of 2–2.5%. Remaining gaps have been filled following methods described by Kieke et al. (2006, 2007).

2.2. Geostrophic transports calculated from inverted echo-sounders In order to study the link between the deep water formation rates and the strength of the subpolar gyre, four PIES (inverted echosounders equipped with pressure sensors) were deployed across the path of the subpolar gyre in August 2006 (Fig. 1). The PIES measure bottom pressure and the vertical travel time of an acoustic signal which is sent from the sea bottom to the surface and received again at the bottom. Their locations are along the western flank of the MAR, sufficiently far away from the very rough small-scale topography at the flanks of the rift valley. The mooring array extends from 471400 N to 521300 N, and the deepest common depth of the four instruments is 3400 dbar, i.e. well below the ridge crest. Judging from the flow field Brambilla and Talley (2006) inferred from drifters, the array is thought to encompass the subpolar gyre. Water crossing the MAR further south most likely stays in the subtropical gyre. The location of the PIES allows the separation of the main known NAC spreading paths through the fracture zones (Charlie-Gibbs, Faraday and Maxwell) at about 531N, 501N and 481N, respectively. Mean daily data from the PIES were retrieved in 2008 by acoustic telemetry, while the instruments remained on the seafloor to complete its scheduled five-year deployment period. Hydrographic data from profiling Argo floats and CTD shipboard measurements are used to calculate the vertical structure of the density variability, the so-called GEM (gravest empirical mode) to convert the round trip travel time measured by the PIES to time series of hydrographic properties and geopotential anomalies (Meinen and Watts, 2000). In total, 64 CTD and 517 Argo profiles, taken between 2006 and 2009 are used to make these calculations. Argo profiles exist all year round and thus cover a wider range of travel times than the CTD profiles, which in general had been taken in summertime. The area from where the profiles were taken is constrained by requesting an unambiguous relation between acoustic travel time and the hydrographic data, and by requesting that all travel times measured by the PIES are covered by the transfer function. The transformation from the observed time series of round trip travel time and bottom pressure into geostrophic transports was done following Meinen and Watts (2000) and Mertens et al. (2009), referenced to 3400 dbar. Argo profiles from the years 2006–2009 were provided and quality-controlled by the German Federal Maritime and Hydrographic Agency (BSH) ¨ following methods described by Bohme and Send (2005), Owens and Wong (2009), and Kieke et al. (2009). The latter also summarized CTD profiles and associated precisions for data prior to 2008. CTD data of 2008 and 2009 have comparable uncertainties in the order of 0.002–0.003 1C for temperature and 0.002 for salinity. The time series of the PIES pressure data suffer from an unknown temporal drift, which is different for each sensor. An exponential–linear function is fitted and subtracted from the pressure time series following Watts and Kontoyiannis (1990). These detrended horizontal bottom pressure differences are used to estimate the geostrophic velocity fluctuations at the reference level of 3400 dbar. The mean absolute velocity at the reference level has to be inferred from other data, i.e. moored current meters and/or combined ADCP/CTD measurements.

2.3. Vessel-mounted and lowered ADCP measurements The velocity distribution at 471N was recorded three times using a self-contained Lowered Acoustic Doppler Profiler (LADCP) system in combination with a CTD probe. The LADCP consisted of two RDI 300 kHz Workhorse Monitor ADCPs operated in synchronized Master-and-Slave mode at a ping rate of 1 Hz and 10 m-depth bins. The raw data were processed as described by Visbeck (2002) yielding full depth profiles of horizontal velocity on a 10 m-depth grid. The

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Sv

Fig. 1. Schematic of the AMOC related circulation in the North Atlantic. The counter-clockwise circulation between 451N and 601N is called the subpolar gyre. NAC: North Atlantic Current, LSW: Labrador Sea Water, ISOW: Iceland Scotland Overflow Water, DSOW: Denmark Strait Overflow Water, C: convection regions. PIES: inverted echosounder with bottom pressure sensor, deployed since August 2006 (at the Mid-Atlantic Ridge) and July 2009 (off Newfoundland). The moorings off Newfoundland at 471N cover the Deep Western Boundary Current and were deployed in July 2009.

30 28 26 24 22 20 18 16 14 12 10 8 6 4 2 0 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004

Fig. 2. Time series of the AMOC volume transport (Sv) at 261N, simulated by the 1/121 resolution FLAME model.

vessel-mounted ADCP data along the full 471N section were taken in 2003 using a 75 kHz vm-ADCP of type Ocean Surveyor on R.V. METEOR, and in 2005 using a 75 kHz ADCP of type Narrowband on R.V. THALASSA. The vertical resolution was 8 m in 2003 and 16 m in 2005. 2.4. Configuration of the 1/121 FLAME model In this study a refined eddy-resolving ocean model of 1/121 resolution is used which is part of the FLAME hierarchy of models for the Atlantic Ocean (Family of Linked Atlantic Model Experiments; Dengg et al., 1999; Eden and B¨oning, 2002). With respect to the subpolar North Atlantic, this model version and its eddy-permitting predecessors with 1/31 resolution of have been frequently used to study mechanisms related to deep water formation (B¨oning et al., 2003; Brandt et al., 2007), eddy variability (Eden and B¨oning, 2002), and mechanisms of AMOC variability (B¨oning et al., 2006; Biastoch et al., 2008). The pathways of deep water export in the Newfoundland Basin and exchanges between the subpolar and subtropical gyre were

studied by Getzlaff et al. (2006) and more recently by Bower et al. (2009) who compared model trajectories to observations from RAFOS float deployments. The simulated domain stretches from 181S to 701N. The simulation uses a horizontal grid of 1/121  1/121cos(j), j being latitude, and 45 vertical levels, with a 10 m-resolution at the surface, smoothly increasing to a maximum layer thickness of 250 below 2250 m. The 10-year spin-up of the model started from a climatology based on a combination of data described by Levitus and Boyer (1994) and Boyer and Levitus (1997). The forcing uses climatological wind stress and heat fluxes as derived from the ECMWF analyses by Barnier et al. (1995). The surface forcing of the following 15-year run is composed of the climatological ECMWF mean (Barnier et al., 1995) by adding interannual anomalies of the NCEP/NCAR reanalyses for the period 1990–2004 (Kalnay et al., 1996). Resulting available model data consists of three-day averages for the period 1990– 2004.The average and the variability of the simulated meridional overturning at 261N (Fig. 2) is comparable to the observed time series inferred from the RAPID/MOCHA array (Cunningham et al., 2007).

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3. Results 3.1. Deep water formation Estimates of time series of the LSW formation rates based on CFC-inventory changes exist since 1997 (Rhein et al., 2002; Kieke et al., 2006, 2007). In Fig. 3, an updated version is presented, including the biennial time series inferred from changes in the CFC inventories between summer 1997 and summer 2005. The estimate for 2003–2005 was about 6% of the formation rate in 1997–1999, indicating a continued significant decrease of the LSW formation rate. As in the previous years, only the lighter LSW mode addressed as ULSW (sY ¼27.68–27.74 kg/m3) was formed. The LSW formation rate time series is compared to an updated transport index (proposed by Curry and McCartney, 2001). Kieke et al. (2007) mentioned already the decline of both time series – the transport

index and the formation rate – and discussed the relation to the NAO index (Fig. 3A). It is well known that the density stratification in the water column in winter and the atmospheric forcing are important in determining the intensity and depth of the convection activity in the Labrador Sea. Mostly, the atmospheric state is reported as the NAO index. The updated versions continue the observed trends of the time series of the last decade. Fig. 3B presents the anomalies of the annual heat loss in the Labrador Sea area. In general, a low NAO index is related to a lower than average heat loss, leading to smaller formation rates (Fig. 3C). Although a quantitative estimate of the formation rate was not possible for 2007, Yashayaev and Loder (2009) concluded from shipboard CTD data and Argo profiles that the convection in the Labrador Sea in 2007 was weak. Convection depths were shallower than 700 m. In winter 2008, the convective overturning extended to 1600 m depth, thus reaching density layers that have

NAO–Index (DJFM)

2 1 0 –1

Annual Heat Loss Anomaly [GJm–2]

–2

1.5 1 0.5 0 –0.5 –1 –1.5

10

75

8

70

7 6

65

5 4

60

3

LSW Formation Rate [Sv]

Baroclinic Transport Index [Sv]

9

2

55

1 50 1991

1993

1995

1997

1999

2001

2003

2005

2007

0 2009

Fig. 3. (A) Annual wintertime index of the North Atlantic Oscillation (NAO) for the years 1990–2009 following Jones et al. (1997). (B) Annual heat loss anomaly (GJ/m2), central Labrador Sea. The mean heat loss from 1974 to 2009 has been subtracted. Surface flux data are from NCEP. (C) Baroclinic volume transport (black line) between Bermuda and the Labrador Sea for the upper 2000 m following Curry and McCartney (2001) for the same period. 3-year low-pass filtered data of this time series is added as a green line. The scale on the right indicates the formation rate of LSW [Sv] shown as red bars each covering 2 years from summer 1997 to summer 2005. Figure is updated from Kieke et al. (2007).

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2500

Layer Thichkness [m]

2000

dLSW:

1500

σΘ = 27.74?27.80 kg/m3

1000

500

uLSW: σΘ = 27.68?27.74 kg/m3

0 1988

1990

1992

1994

1996

1998

2000

2002

2004

2006

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2010

Fig. 4. Temporal evolution of the layer thickness of two modes of Labrador Sea Water in the central Labrador Sea, shown for the years 1988–2009 (updated from Kieke et al., 2006). Blue: upper Labrador Sea Water (ULSW), bounded by the isopycnals sY ¼ 27.68–27.74 kg/m3; Red: deep LSW (in the figure dubbed dLSW), bounded by sY ¼27.74–27.80 kg/m3 (e.g. Rhein et al., 2002; Stramma et al., 2004; Kieke et al., 2006, 2007). A 3-yr filter is applied to smooth the time series (solid lines). Uncertainties are presented as vertical bars.

3.2. Strength of the subpolar gyre

Average density stratification, Labrador Sea 0 200 1998

400

2001 1997

600 2008

Pressure (dbar)

not been renewed since 1994. The intense convection activity was caused by increased oceanic buoyancy loss, associated with below normal air temperatures in this region (Fig. 3B, and Yashayaev and Loder, 2009). In winter 2009, the Labrador Sea fell back into a state of shallower convection depths, which did not reach the deep LSW. As a consequence, the thickness of the deep LSW was reduced again, while the ULSW thickness increased (Fig. 4). Compared to the significant changes observed in former years, the effect of the deep convection observed in winter 2008 on the layer thickness of the two LSW modes was minor (Fig. 4). The thickness increase of the deep LSW observed in summer 2008 was mainly due to the fact that the isopycnal sY ¼27.74 kg/m3 was lifted more strongly than in the preceding and the succeeding year (Fig. 5). The upper boundary of ULSW (sY ¼27.68 kg/m3) in 2008 was at a slightly greater water depth. Therefore, ULSW extended over a more limited depth range than observed in 2007 or 2009.

800

1999

1996

2009

1000

2007

1200 2006

1400 2005

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2003

Up to now, any direct measurement of the strength of the subpolar gyre does not exist. Curry and McCartney (2001) and Kieke et al. (2007) presented a transport index based on calculating the difference in potential energy anomaly between Bermuda and the central Labrador Sea, relative to 2000 dbar. In this way, the resulting index represents the eastward baroclinic mass transport of the upper 2000 m of the water column between these two locations, encompassing transport of both, the subtropical and subpolar gyre. An updated version for the most recent years, expressed as a volume transport, is presented in Fig. 3. In 1997– 2005 a decreasing trend of the LSW formation rate (red bars in Fig. 3) coincides with a lower transport index. H¨akkinen and ¨ Rhines (2004) and Hakkinen et al. (2008) analyzed altimeter data as an independent piece of evidence for a weaker subpolar gyre. From empirical orthogonal functions (EOF) of the geostrophic currents calculated from sea surface height (SSH) measurements they found that the subpolar gyre slowed down in the 1990 s and remained in a weakened state till 2005. They suggested a connection between a weakening of the subpolar gyre and a weakening of the convective forcing in the Labrador Sea and this view is supported by findings of Kieke et al. (2007, and Fig. 3). The transport index, however, represents the transport between the latitude of Bermuda (321N) and the central Labrador

1800 2000 27.65

27.68

27.71

27.74

27.77

27.8

Potential Density σΘ (kg m–3) Fig. 5. Mean density stratification in the central Labrador Sea. Update of figure presented by Stramma et al. (2004) and Rhein et al. (2002). Gray areas highlight the density range addressed as ULSW. Profiles were taken from the central Labrador Sea along the so-called AR7W-section between 511400 W and 521400 W. Chosen CTD profiles were averaged on isopycnal surfaces to generate mean profiles.

Sea at about 581N, i.e. it includes also a part of the subtropical gyre transport. Any co-variability or compensating effects in the strength of the subtropical and subpolar gyres cannot be separated. It is, therefore, highly desirable to have observations that are much more focussed on the subpolar gyre component of the transport. One possible approach is to assess the transport variability of the subpolar gyre as the NAC and underlying water masses cross the MAR and enter the eastern basin of the North Atlantic. Fig. 6 presents the first two years of volume transport variations calculated from the PIES time series of the

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Fig. 6. (A) Volume transport variations between the southernmost (471400 N) and northernmost (531N) PIES. Red: geostrophic transport variation relative to 3400 dbar (27.3 74.2 Sv); blue: transport fluctuations at the reference level (0 7 7.5 Sv); black: both transport time series combined (27.37 7.5 Sv). Positive values represent northeastern flow. (B) combined transport time series (black, dubbed ‘net’) as in top figure, but split into the main water masses. Red: NAC (sy o 27.68), blue: LSW (sy ¼ 27.68–s1500 ¼ 34.72), green: Deep Water (DW, s1500 434.72).

1

Correlation coefficent

0.9 0.8 0.7 0.6 0.5 0.4 0.3 0.2 0.1 0 10°S

0°N

10°N

20°N

30°N

40°N

50°N

60°N

LATITUDE Fig. 7. Correlation between the meridional volume and heat transport in the 1/121 FLAME model. Average over 15 years of model run.

years 2006–2008. The mean transport fluctuation relative to a reference level of 3400 dbar (which is well below the crest of the MAR) is 27.377.5 Sv (Fig. 6A). The dominant part of the fluctuations is caused by the transport variability at the reference level. The NAC (defined as the density range with sY o27.68 kg/m3) covers about 67% of the transport, while the LSW contributes about 28% (Fig. 6B). 3.3. AMOC and the associated heat transport The analysis of the 1/121 FLAME model indicates a close correlation between the AMOC volume and heat transport in

the subtropical Atlantic (Fig. 7). It fits well into the previous range of model solutions (Biastoch et al., 2008) with a weak correlation in the equatorial zone and a correlation maximum of r ¼0.9 (with ‘r’ as the correlation coefficient) around 301N. Northward of 351N the correlation decreases significantly and remains relatively constant in the subpolar gyre at values of about r ¼0.6. Biastoch et al. (2008) attribute the weakened correlation in the subpolar North Atlantic mainly to the increasing importance of the horizontal circulation. Strong flows of eddies in this region might also contribute to a lesser extent to this decorrelation. More details on the underlying mechanisms can be provided when the heat transport is decomposed into the time-mean and

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eddy contributions as well as into zonal variability (gyre component) and mean meridional flow (overturning component). The heat transport calculated from the 1/121 FLAME model has been decomposed at various latitudes between 261N and 541N (Fig. 8). In accordance with results of Kanzow et al. (2009), Fig. 8A reveals the time-mean circulation as the main contributor to the northward heat transport at 261N (1.2 PW), eddy contributions are

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negligible (about  0.1 PW). Although there is increasing eddy variability when going northward into the regions of the Gulf Stream separation and the NAC the contribution of the mean eddy-induced transports remain low (Fig. 8A). The bulk of the heat transport is still given by the time-mean flow. A closer look at the time series of the heat transport components at 471N (Fig. 9A) reveals a much more complicated behavior than

Fig. 8. (A) Time-mean and eddy component of the total meridional heat transport in the 1/121 FLAME model, shown for particular latitudes between 261N and 541N. (B) Contribution of the meridional overturning, and the gyre component to the heat transport in FLAME. Averages over 15 years of model data. The stippled lines denote the standard deviation of the annual means.

Fig. 9. (A) Temporal variability of total meridional heat transport [PW] and associated eddy component across 471N. Time series are calculated from 3-day average values taken from the 1/121 FLAME model. The bold lines denote the respective mean. (B) contributions of the overturning and gyre component [PW] to the meridional heat transport across 471N.

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suspected from Fig. 8A. The mean heat transport of the model at this latitude is 0.6570.17 (one standard deviation) PW, the eddy contribution is calculated as 0.0170.1 PW, i.e. it amounts to 60% of its variability. Decomposition into gyre and overturning components (Fig. 8B) leads to the general conclusion that meridional overturning dominates the heat transport in the subtropical Atlantic, while the gyre component is important for the heat transport in the subpolar gyre. At 471N the contributions of the overturning (0.37 0.1 PW) and gyre component (0.35 70.15 PW) are almost equal. Inspection of the time series (Fig. 9A) indicates that the extreme values of the gyre heat contribution coincide very often with extreme values of the total northward heat transport and thus highlighting the importance of the zonal variability. The correlation between both variables is r ¼0.8, although the variability of the overturning component sometimes counteracts the variations of the gyre component. However, both the gyre and the overturning contributions are important. Thus, the gyre contribution to the heat transport provides the mechanism weakening the correlation between the AMOC and the heat transport in the subpolar gyre. These findings agree with the results of Biastoch et al. (2008). Willebrand et al. (2001) found that near 551N the overturning in density-coordinate models was about 6 Sv larger than the overturning in models calculated in z-coordinates. The authors attributed this effect to a significant contribution of the horizontal gyre circulation to the meridional buoyancy transport. When using density coordinates, the correlation between meridional volume and heat transport is higher than in z-coordinates.

4. Discussion 4.1. Subpolar gyre transport at the MAR The baroclinic transport index from Curry and McCartney (2001) suggests a mean flow of about 50–65 Sv in the upper 2000 m of the water column. As mentioned before, however, this transport estimate encompasses components of the subpolar and the subtropical gyres, which cannot be separated. The biannual mean transport of the subpolar gyre across the MAR referenced to 3400 m depth was calculated to 27.3 Sv, i.e. much lower than the estimates of Curry and McCartney (2001). Using moored current meter data (Schott et al., 2004) as well as current meter and PIES data (Meinen et al., 2000; Meinen, 2001), the top to bottom NAC transport (when subtracting the Mann Eddy) across 421N in the Newfoundland basin (Fig. 1) was estimated to be 90–142 Sv. This is about 3–5 times higher than found at the MAR (Fig. 6). In contrast to that, published estimates of transports west and east of the MAR fit much better to the transports calculated from the Bremen PIES array. Combining hydrography with isopycnal RAFOS floats, Perez-Brunius et al. (2004) estimated the mean eastward flow across 361W (roughly at the MAR) between 471N and 531N to be 21 Sv. Yaremchuk et al. (2001) used geostrophy and a circulation model with a density field evolving with the flow, thereby focussing on the MAR region between 401N–551N and 401W–201W. They reported an inflow across 401W of 51 Sv for the upper 1000 m, and an outflow across the eastern boundary at 201W of 28 Sv. The remaining flow leaves the box to the north and south. Lherminier et al. (2007) came up with a NAC transport of 21–28 Sv between 451N and 541N. This was inferred from two one-time hydrographic sections running almost parallel to the Bremen PIES array but on the eastern side of the MAR. Paillet and Mercier (1997) and Cunningham (2000) estimated a somewhat higher transport of 35 Sv across the MAR between 401N and 541N. The calculated LSW inflow into the eastern Atlantic of 7.8 Sv (Fig. 6) is compatible to the results found by

inverse modeling (6.3 Sv; Paillet et al., 1998), but higher than the 2.6–3.1 Sv LSW crossing the MAR between 451N and 541N reported by Lherminier et al. (2007) based on two hydrographic sections. One has to bear in mind that the transport estimates presented here (Fig. 6) are referenced to 3400 dbar, while the other estimates, except the transport index (Fig. 3), are thought to be absolute transports. The reference level of 3400 m is located well below the ridge, leading to the assumption that the total transport might be not much different from the estimates reported here. Read et al. (2010) used a one-time hydrographic survey from 491N to 541N at the MAR. They found the geostrophic shear between 1000 and 2500 m quite similar to the across-track velocity measured with the LADCP. The best match between both data sets was achieved with velocities around 73 cm/s at the deepest common level, and the authors concluded that the geostrophic transport relative to the bottom roughly equals the total transport. The discrepancy between the large northward NAC transport at 421N and the much smaller transport across the MAR into the eastern Atlantic could have several reasons. Surface drifters deployed from 1990 to 2002 in the subtropical Atlantic showed that a high percentage of the floats passing 421N cross the MAR south of 471N and remain in the subtropical gyre (Brambilla and Talley, 2006). The trajectories of RAFOS floats deployed in the NAC in 1993–1995 indicated a NAC which closely followed the continental slope till 511N before turning eastwards (Rossby, 1996). Some of the floats remained in the Newfoundland Basin, Irminger Sea, and Labrador Sea and did not cross the MAR during their lifetime (Rossby and Prater, 2005). The main NAC crossing pathway at that time period was the Faraday Fracture zone at 511N, but Bower and von Appen (2008) showed that this is not always the case. In other years, several different branching modes occurred, with the main pathway fluctuating between 531N and ¨ 491N. Hakkinen and Rhines (2009) analyzed surface drifter tracks form the time period 1990 to 2007. Since 2001 drifters deployed south of 451N–471N within the warm water domain reached farther to the north in the Northeast Atlantic instead of turning to the southeast like in the years before. The inferred shift of the subpolar front to the northwest, i.e. towards the Charlie-Gibbs Fracture Zone, was not observed in our time series, the main NAC path was more to the south. The fraction of the NAC staying in the western Atlantic might be dependent on the latitude of the main crossing path. Simultaneous time series of the meridional NAC transport in the Newfoundland Basin (for instance at 471N) and of the gyre transport crossing the MAR would be very helpful to address this question.

4.2. Meridional heat and volume transport variability Inverse model calculations of heat and mass transports at 431N from five hydrographic sections between 1957 and 2000 led to the conclusion that the MOC volume transport at that latitude does show significant interannual variability of 13–20 Sv (Koltermann et al., 1999; Lorbacher and Koltermann, 2000). The same hydrographic data combined with other constraints (Lumpkin et al., 2008) resulted in nearly constant volume transports in all the five repeats of 16– 18 Sv, showing that the results are strongly dependent on the assumed constraints. Lherminier et al. (2007) used combined hydrographic and LADCP data from two sections between Greenland and Portugal taken 2002 and 1997 and calculated 16.9 and 19.2 Sv, respectively. The 1/121 FLAME model exhibits a 15 year (1990–2004) mean overturning at 471N in the subpolar North Atlantic of 10 Sv, with a standard deviation of 3.2 Sv and short-term variations of up to 12 Sv, indicating that transport variability observed by one-time surveys are unable to separate long-term trends from short-term

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variability. The maximum overturning in this model of 16 Sv is located further south (Biastoch et al., 2008). When analyzing high-resolution (1/121) North Atlantic (FLAME) and global ocean-ice (ORCA) models, Biastoch et al. (2008) examined the cause for interannual to decadal variability of the AMOC. The AMOC fluctuations were linked to the variability of the deep water formation in the Labrador Sea, but this signal is effectively masked by stronger high-frequency variability related to wind forcing, and north

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of 351N also by internally induced (eddy) fluctuations. Getzlaff et al. (2005) and B¨oning et al. (2006) studied the thermohaline AMOC signal in isolation in the FLAME model. It was found to be of the order of 2 Sv north of 401N. In the FLAME and ORCA models, the maximum of these anomalies is located at the southern exit of the subpolar gyre near 451N (Biastoch et al., 2008). The standard deviation of the monthly AMOC time series, however, is significantly larger (3 Sv) than the thermohaline signal (Biastoch et al., 2008). The

Fig. 10. (A) Meridional velocity [cm/s] along 471N averaged over 15 years of 1/121 FLAME data, red northward, blue southward velocities. Black lines: standard deviation in cm/s. (B) 15 year mean temperature distribution [1C] for the same section. Black lines denote the associated standard deviation (1C).

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35.51W mean and variability of the velocity field are much less structured and less intense. The variability in the circulation is mirrored in the fluctuations of the temperature distribution (Fig. 10B). The strongest variability in the western basin is caused by the meandering of the front between the cold southward boundary current and northward flowing warm NAC water. In the eastern Atlantic the variability is smaller and dominated by the seasonal signal in the mixed layer. The model does not show the existence of any eastern boundary current, either in the strength of the velocity or in enhanced variability of the current or of the temperature field. In conclusion, the important circulation components dominating the heat transport and its variability in the 1/ 121 FLAME model are located in the western basin. Outside the western boundary current, the spatial scales of the velocity and temperature features in the FLAME model are of the order of 120–160 km (radii 60–80 km). Kieke et al. (2009) compared time series of temperature and salinity recorded at 1500 m by moorings at the western flank of the MAR, with Argo profiles measured nearby. Profiles less than 80 km away from the mooring location showed the best correlation with the records from the mooring site. At greater distances the Argo data were in general not correlated with the mooring records. Given the high level of temporal variability in the moored instruments, a distance of 80 km could be taken as a rough measure of a spatial correlation scale. In summer 2005, the subpolar North Atlantic was surveyed along roughly 471N with 61 XBTs and 39 CTD profiles, i.e. the horizontal resolution of the temperature field was about 30–36 km for the western and the eastern basin, respectively. The spatial scales found on this section are about 80 km (Fig. 11). Two velocity sections derived by 75-kHz vessel-mounted ADCPs exist that were taken in the upper 600–700 m along 471N in 2003 and 2005 (Fig. 12). Features identified from the two snapshots of meridional velocities are similar, with the highest velocities in the western basin west of 271W and weaker signals in the eastern basin. As in the model, an eastern boundary current is not recognizable in the observations (Fig. 12B). The spatial scales are similar to the ones present in the temperature field (Figs. 11 and 12). The eddy radius was estimated in a larger area of the subpolar North Atlantic by combining altimeter data with hydrography (Fig. 13). Only those eddies which have been sampled and validated with CTD stations are presented. South of about 501N, the radius of the eddies is larger than 70 km, while further north the radius decreases as

standard deviation increases up to 15 Sv when taking daily means. Each one-time survey contains this kind of variability making it very difficult if not impossible to extract the desired long-term trends. Thus a continuous time series for the AMOC volume and heat transport is necessary.

4.3. Elements of an AMOC array in the subpolar gyre One such AMOC array is already in place in the subtropical Atlantic at 261N (e.g. Cunningham et al., 2007). The RAPID/MOCHA array consists of cable measurements across the Florida Strait, current meter and T/S moorings at the eastern and western boundaries, as well as PIES. The AMOC transport is separated into the Florida Current, the meridional Ekman transport, and the internal transport, which is decomposed into three observable contributions. The transport at the western boundary is inferred from direct velocity measurements, the geostrophic internal transport is calculated from the density difference between the eastern and the western boundary relative to a reference level. Zonal differences of ocean bottom pressure fluctuations allow the computation of the meridional velocity fluctuations at the reference level. Cunningham et al. (2007) report a year-long average AMOC transport of 18.775.6 Sv, with a large short-term variability ranging from 4.4 to 35.3 Sv over a year, confirming the variability found in high-resolution models (e.g. Biastoch et al., 2008). The composition of the meridional volume and heat transport in the subpolar North Atlantic requires to monitor temperature and velocity fields simultaneously. The FLAME model is used to evaluate how an observing array to monitor the AMOC volume and heat transport in the subpolar gyre could be designed. To investigate in more detail the regions where high resolution in the observations might be needed, Fig. 10A shows the 15-year model mean of the velocity field and its standard deviations along the 471N section. The flow in the Flemish Pass at 471W is towards the south with mean speeds exceeding 40 cm/s. The narrow southward western boundary current is visible between 441W and 431W. Next to it a strong, surface enhanced northward flow associated with the NAC extends as far east as 401W. A weaker southward branch reaches from 401W to 381W and a northward component from 381W to 35.51W. The standard deviation is high in any of the mentioned areas and is found throughout the whole water column with highest values near the surface. East of

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expected. The results confirm an earlier study by Krauss et al. (1990), who inferred from altimeter and surface drifter data eddy length scales between 60 and 95 km near 471N, and around 50 km further north. Similar eddy length scales were also found by Eden (2007) in altimeter data and in the 1/121 model. Both observations and FLAME model results agree on the magnitude and the northward decrease of the spatial scales of the flow and temperature in the subpolar North Atlantic. Since at least some of these features in the western basins have to be resolved to obtain reliable estimates of volume and heat transport, regions with smaller eddy length scales probably are quite cost-intensive and require large instrumental resources. The area south of 501N is therefore more preferable to locations further to the north. When focusing on this region volume and heat transport through the Flemish Pass have to be considered. In the model, the mean Flemish Pass flow is southward from top to bottom and higher than 20 cm/s (Fig. 10). One section of the few existing ADCP measurements from July 2009 shows a more complex characteristic than in the model, with colder and fresher water flowing southward on the western side and a northward flow of warmer and more saline water on the eastern flank of the Pass (Fig. 14). The latter could be part of the anticyclonic circulation often found around the shallow Flemish Cap, separating Flemish Pass from the open ocean farther east (Gil et al., 2004). In the Pass, the flow is mainly barotropic (Fig. 14), and thus does not require many instruments to be deployed, but at least two mooring sites are needed to resolve the flow and temperature field. Owing to the relatively steep topography, the western boundary current at 471N is narrow, so that three moorings with T/S sensors and acoustic current meters are sufficient to resolve the boundary current. In July 2009, three moorings and an additional PIES were deployed at 471N along the eastern flank of Flemish Cap

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(Fig. 1). While the mooring array was designed to capture the southward export, the adjacent PIES combined with the PIES array parallel to the MAR provides the integral transport variability in the interior Newfoundland Basin. Judging from Figs. 10 to 15, a spatial resolution of about 80 km would be needed here to measure the meridional volume and heat transport at this latitude. In 2010, an additional PIES was deployed at about 391W, separating the main NAC import path in the west from

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the recirculation zone to the east. Judging from the mean model velocity field and from the three LADCP snapshots (Fig. 15), the two western PIES encompass the northward NAC flow across 471N, providing sufficient data to compare the NAC transport fluctuations there and across the MAR. The LADCP snapshots from the Newfoundland Basin resemble quite closely the mean current field of the FLAME model (Fig. 10): east of the narrow boundary current, a strong northward band west of 40–391W (the main NAC inflow) is followed by a southward flowing band, and the structures further east are weaker than in the western Newfoundland Basin. The many similarities between model and observations make the model an effective tool for planning the observation system.

5. Conclusions Separating natural heat and volume transport variability from the trend expected in a global warming scenario requires long observational time series and thus an array, which could be efficiently maintained over several decades. Such an array in the interior could consist of C-PIES (PIES equipped with a current meter). A sufficiently close spacing of such instruments will allow measurement of the baroclinic and barotropic velocity fluctuations from the acoustic travel time and bottom pressure measurements and the mean velocity field at the depth of the current meter.

Measuring these barotropic flows in the subpolar North Atlantic has been shown to be a critical component of an AMOC observing system (Baehr et al., 2004). Observations and model show larger velocities and variability in the western Atlantic, suggesting that variability of the volume transport is dominated by contributions from the western basin. Thus, the horizontal resolution of the C-PIES array can be coarser in the eastern basin. Variability in the eastern boundary current or temperature are comparable to that in the eastern interior, so it might be acceptable to deploy C-PIES at the continental slope instead of conventional moorings. Altogether, about 30–35 C-PIES would be sufficient to cover the 471N section east of the western boundary. The data of all instruments could be retrieved by acoustic telemetry in less than 180 hours ship-time per year. Combining the C-PIES readout with a CTD survey and the recovery and redeployment of the moorings at the western boundary and in Flemish Pass requires about 4 weeks of a research vessel per year, which is comparable to the requirements of other oceanographic projects. A prerequisite for the successful use of C-PIES in the future, however, is the continuation of the Argo program. It is expected that the T/S characteristic of the various water masses might change with time. Sufficient T/S profiles year round are therefore needed to be able to monitor these changes. The continuation of the LSW formation rate time series and the measurement of the DWBC transport also remains a necessity to understand the relation between these three major components of the Atlantic climate relevant circulation.

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Hamburg. The Argo data used for calculating GEMs were collected and made freely available by the International Argo Project and the national Argo programs that contribute to it (http://www. argo.ucsd.edu, http://argo.jcommops.org). Argo is a pilot program of the Global Ocean Observing System. Time series of annual heat loss anomalies in the Labrador Sea were computed from the Reanalysis Project of the U.S. National Center for Environmental Prediction (NCEP, http://www.cdc.noaa.gov). Further thanks go to the team of the Bermuda Atlantic Time-Series Study (BATS) for making hydrographic data available that were used to calculate the baroclinic transport index. Data were received via the onlinedata extraction interface at http://bats.bios.edu.

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Acknowledgments This work is part of the national cooperative project ‘Nordatlantik’, funded by the German Ministry for Education and Research (BMBF), Grants 03F0443C and 03F0605C (both to M. Rhein). The position of A. Roessler was funded by the German Cluster of Excellence (MARUM). The German Research Foundation (DFG) supported the THALASSA cruise, providing charter and other funds. We acknowledge the contributions of J. Dengg, R. Redler, J.-O. Beismann, C. Eden, and L. Czechel to the FLAME development and integration. The computations were performed at DKRZ,

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