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  • View in gallery

    Map of the locations of the moorings described in this report (Y–G), of two moorings (252 and 253) described in the text (Clarke 1984), and of two moorings to the north (ACM8); see text. The bathymetry is shown in meters.

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    Cross section of the moorings, showing locations of the instruments. Mean depths of the 3°C isotherm and of the density surfaces σθ = 27.85 and 27.8 kg m−3 are also shown.

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    (top) Potential temperature (θ) on the mooring line derived from CTD casts made during the deployment of the moorings (4–6 Sep 2005) showing the cold layer on the seabed, and (bottom) the corresponding salinity showing the saline layer above the cold layer. Both are components of the DWBC.

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    Salinity (upper trace) and potential temperature (lower trace) for the entire record of RCM h4, i.e., the instrument on mooring H at depth 2365 m, 53 m above the seabed. The inset panel shows in detail the first 2.5 days of the salinity record.

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    Record-mean current vectors at each IFREMER and NOCS mooring: the current increases almost everywhere toward the bottom and closely parallels the bathymetry (depths in m).

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    The character of the current in the core of the DWBC (h4, main panel), and on its edge (e3, inset) as seen in virtual displacement plots, both approximately 50 m above the seabed. Elapsed time is marked every 10 days, with every 50 days numbered.

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    The DWBC speed fluctuations on four instruments (denoted by color) on mooring B in the direction 032°: the means have been removed, and the vertical coherence of the fluctuations measured at four locations over the depth range 1980–2600 m is striking.

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    Mean value of the transport per unit width, on each mooring, colder than 3°C (long dash) and denser than 27.85 kg m−3 (short dash), with estimates of the likely errors of the former.

  • View in gallery

    Time series lasting 284 days of the 12-h average DWBC total transport colder than 3°C. The superimposed circles are 20-day-average values.

  • View in gallery

    Spectral analysis of the total transport colder than 3°C displayed in variance-preserving manner and similar to figures shown for transports at upstream locations (Saunders 2001). The logarithm of the frequency (cpd) (logfreq, primary x axis) is plotted, and the equivalent period (days) is shown on the secondary x axis.

  • View in gallery

    Contributions to the DWBC transport (Sv) colder than 3°C from each mooring for (a) the record average and for extreme conditions: two cases of (b),(c) high and (d),(e) low transport. Date range and total transport are shown for each panel.

  • View in gallery

    Potential temperature vs salinity for CTD stations on the mooring line: data from the NOCS deployment cruise (298, 4–6 Sep 2005, black) and the NOCS recovery cruise (309, 22–31 Aug 2006, red). The warming and salinity increase below 3°C is evident. The three density lines are described in the text.

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    CTD station positions for VEINS cruises. The data shown are for the 1997 cruise and were obtained from the NODC WOCE archive.

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    Year-long current vectors on the Cape Farewell section approximately 110 m above the seabed. The black vectors refer to the data reported here (moorings Y–C), and the gray vectors to data from 1995 and 1996. See text for source (WOCE).

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    Core properties in the overflow on the density surface σ2 = 37.12 kg m−3 derived from the NODC WOCE data archive, the data of Holfort and Albrecht (2007), and after 2004 recent Discovery cruises.

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The Deep Western Boundary Current at Cape Farewell: Results from a Moored Current Meter Array

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  • 1 National Oceanography Centre, Southampton, Southampton, United Kingdom
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Abstract

An analysis is made of data from 30 Aanderaa recording current meters (RCMs) set on nine moorings located east of Cape Farewell, the southern tip of Greenland. The purpose of the measurements was to allow for the estimation of transport in the deep western boundary current (DWBC) below a depth of about 1500 m. The records commenced in September 2005 and lasted from 9.5 to 11.5 months. After calibration of the raw data, 12-h averages of temperature and current were derived and the latter employed to estimate the flow across and along the array direction. The 9.5-month average transport of water colder than 3°C was found to be 7.8 Sv (1 Sv ≡ 1 × 106 m3 s−1) with a standard error of 0.8 Sv. For water denser than σθ = 27.85 kg m−3, the transport is calculated as 4.5 Sv. Whether either of these values is significantly different from comparable measurements made 500 km upstream cannot be determined. In marked contrast, for σθ > 27.8 kg m−3, the transport is estimated as only 9.0 Sv, smaller than the widely accepted value of 13 Sv for nearby measurements made in 1978. A reevaluation of the calculations and assumptions made then allows one to determine the uncertainty of the earlier estimate and thereby conclude that the difference between the previous and present measurements is significant, that is, that the transport has decreased between 1978 and 2005–06. A weakening of the transport during the 9.5-month period is also observed, along with a warming and an increase in salinity in the core of the DWBC. These latter changes are shown to be consistent with interannual variability rather than a long-term trend.

Corresponding author address: Dr. Sheldon Bacon, National Oceanography Centre, Southampton, University of Southampton, Waterfront Campus, European Way, Southampton SO14 3ZH, United Kingdom. Email: s.bacon@noc.soton.ac.uk

Abstract

An analysis is made of data from 30 Aanderaa recording current meters (RCMs) set on nine moorings located east of Cape Farewell, the southern tip of Greenland. The purpose of the measurements was to allow for the estimation of transport in the deep western boundary current (DWBC) below a depth of about 1500 m. The records commenced in September 2005 and lasted from 9.5 to 11.5 months. After calibration of the raw data, 12-h averages of temperature and current were derived and the latter employed to estimate the flow across and along the array direction. The 9.5-month average transport of water colder than 3°C was found to be 7.8 Sv (1 Sv ≡ 1 × 106 m3 s−1) with a standard error of 0.8 Sv. For water denser than σθ = 27.85 kg m−3, the transport is calculated as 4.5 Sv. Whether either of these values is significantly different from comparable measurements made 500 km upstream cannot be determined. In marked contrast, for σθ > 27.8 kg m−3, the transport is estimated as only 9.0 Sv, smaller than the widely accepted value of 13 Sv for nearby measurements made in 1978. A reevaluation of the calculations and assumptions made then allows one to determine the uncertainty of the earlier estimate and thereby conclude that the difference between the previous and present measurements is significant, that is, that the transport has decreased between 1978 and 2005–06. A weakening of the transport during the 9.5-month period is also observed, along with a warming and an increase in salinity in the core of the DWBC. These latter changes are shown to be consistent with interannual variability rather than a long-term trend.

Corresponding author address: Dr. Sheldon Bacon, National Oceanography Centre, Southampton, University of Southampton, Waterfront Campus, European Way, Southampton SO14 3ZH, United Kingdom. Email: s.bacon@noc.soton.ac.uk

1. Introduction

The deeper, southgoing limb of the Atlantic Meridional Overturning Circulation comprises North Atlantic Deep Water (NADW), which in turn is made of upper (uNADW) and lower (lNADW) components. The uNADW is largely sourced from Labrador Sea Water, and lNADW derives from dense waters of the Nordic Seas and Arctic origin that overflow the Greenland–Scotland Ridge (e.g., Bryden et al. 2005; Yashayaev and Dickson 2008). The most southerly point at which the Greenland–Scotland overflows can be observed before they enter the Labrador Sea is off Cape Farewell, the southern tip of Greenland. However the magnitude and variability of the transport of the deep western boundary current (DWBC), which contains the overflow waters east of Cape Farewell, is poorly understood. On the one hand, Clarke (1984) describes the derivation of the presently widely accepted value of 13 Sv (1 Sv = 1 × 106 m3 s−1), based on a combination of a hydrographic section with a 60-day deployment of a small array of current meters; on the other hand, smaller transports have been calculated from hydrographic “snapshots” (e.g., Bacon 1997; Lherminier et al. 2007). To attempt to resolve this apparent contradiction, the National Oceanography Centre, Southampton (NOCS), deployed an array of moored current meters east of Cape Farewell for a year, as a contribution to the U.K. Natural Environment Research Council’s Rapid Climate Change Programme.

The NOCS mooring array was deployed during RRS Discovery cruise 298 (D298, Bacon 2006a) in September 2005 and recovered during RRS Discovery cruise 309 (D309, Bacon 2006b) in August 2006. D298 also comprised a hydrographic survey of the region around Cape Farewell, and some results from this survey have been reported in Holliday et al. (2007), describing a partial retroflection of the East Greenland Current; in Lauderdale et al. (2008), describing the inferred turbulent mixing in the DWBC; and in Holliday et al. (2009), who examine the three-dimensional circulation around Cape Farewell employing currents measured by lowered acoustic Doppler current profilers (LADCPs).

Seven NOCS moorings with 27 Aanderaa current meters were deployed from 4 to 6 September 2005 and recovered from 22 to 31 August 2006. Of the 27, 2 failed to yield data, while 25 (9 RCM 7/8 and 16 RCM 11) yielded satisfactory current and temperature records. Moorings were identified by a single letter (A–H) and instruments by a digit (1–5), the latter numbered from the top down; thus, b1 is the shallowest instrument on mooring B and b5 that nearest the seabed.

In June 2004 (15 months earlier than the start of the NOCS experiment), personnel from the Institut Français de Recherche pour l’Exploitation de la Mer (IFREMER, Brest, France) deployed moorings inshore of the array described above. They were investigating the East Greenland Current, a surface-intensified western boundary current. These moorings were recovered 19–23 June 2006 and therefore end 2 months before the NOCS records, but there is an overlapping period of 9½ months. The calibration of the IFREMER data is not described here, but our colleagues have generously made available data from the deepest instruments (RCM 8) on two of their moorings and we have included these in our analysis.

The IFREMER moorings were identified by an identical set of letters to ours, so to avoid confusion we have renamed their mooring B as mooring Y here and their mooring A as mooring Z. The locations of all moorings are shown in Fig. 1 and the positions and depths of the moorings and instruments are listed in Table 1. Figure 2 shows a cross section of the moorings, displaying the vertical disposition of the current meters. Figure 3 is the CTD section on the deployment cruise (D298) revealing both the cold bottom boundary component and the saltier and warmer overlying component, whose partial and combined transport is to be estimated from the data.

The structure of the remainder of this paper is as follows. In section 2 we describe the calibration and processing of the data from the NOCS instruments. In section 3 we describe the means, variability, and trends in the computed DWBC transports. Section 4 presents a discussion of the results. Readers uninterested in the details of the calibration of the instruments may omit the following section and pass directly to section 3.

2. Calibration and data processing

A considerable volume of calibration data existed for the 25 NOCS instruments. On stations during the recovery cruise (D309), between two and seven instruments in rapid recording mode were strapped to the CTD frame. Where stops were made on a CTD cast to collect water samples, comparisons were made between the pressure, temperature, and conductivity of the CTD and corresponding measurements from the current meters. These data were analyzed by S. Alderson (NOCS) and provided a set of on-cruise calibration coefficients for all NOCS RCMs. The RCM 7/8 instruments also had precruise calibrations carried out at NOCS and the RCM 11 instruments also had factory calibrations, which did not include conductivity. We employed both sets of calibration coefficients in converting raw data to physical measurements. Eventually, we needed to make a choice and selected the precruise calibrations for pressure, as they more closely agreed with the instrument depth/pressure, and the on-cruise calibrations for temperature as they provided a homogeneous dataset. The fate of conductivity/salinity will be described later. Current speed and direction had standard calibrations from the manufacturer. On the issue of timing, since the RCM 7/8 recorded throughout each hour (20 samples at 3-min intervals), the midinterval time was selected for each measurement. In contrast, the RCM 11 recorded a burst of data just before each hour and each half hour, so these were the times assigned to the measurements. In general, the number of data cycles between start time and stop time agreed with the elapsed time within one or two data cycles, but for four instruments this was not the case. For three of these instruments (a3, b5, and c3), the first and last record in the water were easily identified, but the number of intervening cycles was fewer than expected. For record c2 the recovery data cycles were missing.

Plots of temperature, pressure, and salinity were made for each instrument. It was evident that all temperature and pressure records were good, save for the pressure records from d4 (noisy) and a4 (showed drift). Because the moorings were constructed to be stiff/taut, the knockdown was only intermittent, even in the core of the current. Nevertheless, knockdown was useful because it occurred simultaneously on all of the records on a mooring and could be used to identify missing data for the four instruments indicated above. Absent data was thus introduced into record a3 (four data gaps of 3, 5, 10, and 20 days), into b5 (one gap of 33.5 days), into c3 (one gap of 5.5 days), and c2 (one gap of 2.5 days immediately before recovery)—the number of data cycles augmented accordingly.

Conductivity, and therefore salinity, measurements (present on all instruments) were disappointing. Figure 4 shows the temperature and salinity record from instrument h4, one of the quietest and cleanest salinity records. The upward drift of salinity of ∼0.1 for the first few days illustrated here occurred in all of the salinity records from the 16 RCM 11 instruments and in some was twice as large. We believe this behavior to be entirely instrumental. Of the nine RCM 7/8 instruments only three seemed at all plausible and between four and six measurements were very noisy, obscuring any signal. Accordingly, we have not felt it worthwhile to pursue the salinity measurements further. A final step in the calibration procedure was to correct the current direction from magnetic to true (deviations varied from 22° to 24°W). Figure 5 shows how the record mean currents generally followed the bathymetry and Fig. 6 shows their differing nature: h4 is in the core of the current and e3 in the offshore edge.

Most of the results shown in subsequent figures and those described in the body of this paper were produced from a 12-h-averaged dataset centered on midday and midnight. A similar averaging was undertaken for the 1-h data supplied from IFREMER for moorings Y and Z. For the NOCS records a common dataset was derived containing 701 values commencing on 6 September 2005 at 1200 UTC and ending on 22 August 2006 at 1200 UTC: the IFREMER data that we employed here commenced at the same date and time but ended 16 June 2006. Temperature was converted to potential temperature making use of the CTD data closest to each mooring from both the deployment and recovery cruises. Temperature minus potential temperature is almost entirely a function of pressure (depth) and ranged from 0.093°C on instrument z3 to 0.240°C on g3. The line of moorings was oriented in the direction 122° to 302°, and for each record the variables east and north were replaced by components in the directions 032° (crossing the mooring line) and 122° (along the mooring line). The last step in the data preparation was to fill the gaps in the “short” records. It was noted that velocity fluctuations seen on a mooring were essentially independent of depth, that is, were barotropic. See the example from mooring B in Fig. 7. Accordingly, the gaps in the short record were filled from the fluctuations of the nearest instrument added to the mean of the short record. This procedure was employed for the rotated velocity components and for the potential temperature. A summary of all final estimates is found in Table 2.

3. Transports

a. Means

Our aim is to construct DWBC transport estimates for comparison purposes and as a benchmark for numerical model calculations. Accordingly, we shall employ limits to the definition of the transport involving both potential temperature (3°C) and the density parameter σθ (with values of both 27.8 and 27.85 kg m−3). For the latter definitions, we are hampered by the inaccuracy of or lack of salinity measurements on the RCMs, and so must rely on CTD stations made during deployment and recovery. These three limits are conventional in the literature (e.g., Saunders 2001; Dickson et al. 2008). The lower density (27.8 kg m−3) separates Greenland–Scotland overflow waters from overlying waters. The higher density (27.85 kg m−3) separates Iceland–Scotland overflow water from the deeper-lying Denmark Strait overflow water. The temperature (3°C) is used for comparison with similar published estimates and, typically, lies deeper than σθ = 27.8 kg m−3. These waters are described further below and in section 4.

For transport colder than 3°C we adopted the following method. On the NOCS moorings (A–G) the current meter with mean potential temperature nearest 3°C was found to be the shallowest. At this level the gradient of potential temperature with height was obtained from nearby CTD stations and at each time step the difference between the observed potential temperature and 3°C was converted into a height difference. When added to the depth of the instrument, the depth of the 3°C isotherm was determined and hence a time series constructed. For the IFREMER moorings (Y and Z) the depth of the 3°C isotherm could be interpolated at each time step from the RCM records themselves. The mean, that is, 9½-month depth) of this isotherm is shown in Fig. 2 along with the mean depth of the density surface, σθ = 27.8 and 27.85 kg m−3, derived from NOCS deployment and recovery cruises.

By integrating the crossing components (direction 032°) on a mooring with respect to depth we have constructed the transport time series colder than 3°C per unit distance (in the direction 122°), along with the mean value, variance, and error of the mean. The standard error ε is derived from the decorrelation time scale of the series τ and the record length T (to provide the number of independent estimates of transport), and the variance σ2 according to the expression
i1520-0485-40-4-815-eq1
The decorrelation time scale, obtained from the zero crossing of the lagged autocorrelation function, increases with distance downslope from about 4 days at Y to 16 days at G. The standard error increases in a similar manner and is shown along with the mean transport at each mooring in Fig. 8. To obtain the entire transport in the DWBC, integration is then made along the mooring line, that is, between the moorings. The mooring G data, with a mean northward recirculation, has been excluded from the along-section integration.

For transport colder than 3°C, the 9½-month mean (the duration limited by the June recovery of the IFREMER instruments) is found to be 7.8 Sv with ε = 0.8 Sv. For water colder than 3°C the transport-weighted potential temperature is found as 2.36°C with a probable error of 0.05°C. For water denser than 27.85 kg m−3, a transport value of 4.5 Sv is obtained. The error is difficult to quantify because, although we have a time series for the currents at each mooring location, we have only a pair of values (from the deployment and recovery cruise CTD data) for the thickness of the dense water there. A similar comment applies for transport denser than 27.8 kg m−3 for which the value is determined as 9.0 Sv. This density surface lies between 100 and 250 m shallower than the 3°C isotherm, contributing an additional 1.2 Sv on average above the transport below 3°C, so we believe its standard error to be 1 Sv. The basis for these values is to be found in Table 3. We shall comment on these results in the discussion section.

We note that the DWBC has two clearly separated “cores”: a “transport” core, associated with the maximum current speeds seen at moorings Z–H, where the water depth is ∼1900–2450 m, and a “property” core, associated with the DWBC deep temperature minimum in the vicinity of moorings C and D, where the water depth is 2900–3000 m. The property core may have a secondary transport maximum (Fig. 8). The two cores are separated by a distance of ∼100 km. Dynamical analysis of this observation will be pursued elsewhere.

b. Variability

The time series for transport colder than 3°C is shown in Fig. 9 and the standard deviation (σ) of the series overall is 2.9 Sv. The minimum transport in the direction 212° is 0.9 Sv and occurs in May 2006; the maximum is 16.4 Sv and occurs during November 2005. Variability is shown on a wide range of time scales, as revealed by the 20-day averages, Fig. 9, and by a spectral analysis of the time series, Fig. 10. Most of the variability is in periods from 10 to 50 days.

What is the structure of this variability? Correlation analysis shows that, in the core of the flow, adjacent mooring transports are only very weakly correlated; that is, the eddying fluctuations are independent. Offshore of the core some slightly stronger negative correlations are found, suggesting eddying structures can be slightly larger than the mooring separation (∼30 km). Employing rotary spectral analysis, we have examined the eddying motions to determine if they are either predominantly cyclonic or anticyclonic, but they are found to have equal frequency within the overflow. To illustrate the unresolved spatial nature of the fluctuations we show in Fig. 11 the contributions to the total transport on a few occasions, encompassing both extreme and average conditions.

c. Trends

Not only is the DWBC transport highly variable at periods less than 50 days but it decreases, as Fig. 9 shows: it is stronger at the start than at the end of the record. To quantify the change we have calculated the transport for the first 60 days and for the last 60 days and these values are 9.4 ± 1.9 Sv and 6.1 ± 2.2 Sv (colder than 3°C) and are shown in Table 4. The potential temperature also changes over the record, as seen in a temperature–salinity plot of CTD data on the mooring line from the predeployment, D298, and the recovery, D309, cruises (see Fig. 12). Below 3°C evidently, the water column warmed and became more saline; the warming can also be seen in Fig. 4. Thus, we have included in Table 4 temperatures and changes from the near-bottom instrument on each mooring: the latter values range from −0.08° to +0.51°C, with only one of the nine instruments indicating a cooling and the remainder indicating a warming. These changes are considered in section 4.d below.

4. Discussion

To describe present understanding of the circulation of the northern overflows, we quote from Quadfasel and Kase (2007): “Dense Nordic waters enter the North Atlantic through passages in the Greenland–Scotland Ridge at a mean rate of 6 Sv. Subsequent entrainment of ambient water into the sinking plumes downstream of the sills approximately doubles this flux. Decade long observations show these fluxes to be stable with no discernable trends.”

The overflow sources are Denmark Strait, west of Iceland, and, to the east of Iceland, principally the Faroe Bank Channel. They are approximately of equal magnitude. The overflow east of Iceland proceeds down the eastern flank of the Mid-Atlantic Ridge and makes its way into the western basin via the Charlie Gibbs Fracture Zone. South of Denmark Strait the two combine, with the eastern waters warmer and more saline and lying above the colder, fresher waters arriving over a shorter path from Denmark Strait; here seen in both Figs. 3 and 12.

a. Transport colder than 3°C

On the Ammassalik (formerly known as Angmagssalik) section 550 km south of the Denmark Strait sill (section 3 in Fig. 13), an array of current meter measurements made during 1987–90 by Dickson and Brown (1994) yielded a steady transport colder than 3°C of 9 Sv with ε = 0.25 Sv and σ = 3.0 Sv (Saunders 2001). While these upstream values are a little larger than, but similar to, our Cape Farewell mean of 7.8 Sv (ε = 0.8 Sv, σ = 2.7 Sv), the transport-weighted potential temperature is significantly different, namely, 1.78 ± 0.03°C at Ammassalik and 2.36 ± 0.05°C at Cape Farewell, indicating considerable mixing and entrainment/detrainment.

Support for continuing mixing within the DWBC down the entire length of the east Greenland margin between Denmark Strait and Cape Farewell, and farther downstream into the Labrador Sea, is given by Lauderdale et al. (2008), who employ some of the same CTD data (D298 and others) as used here. They show that “moderate” mixing (10−4 m2 s−1) over the greater (1000 km) distance to Cape Farewell is as important as strong mixing (10−3 m2 s−1) near (within 100 km) the Denmark Strait sill. Part of their evidence is provided by the downstream-decreasing gradient in the DWBC core density: over the 100 km near the sill, a decrease of 0.1 kg m−3 is followed by a further decrease of 0.1 kg m−3 over the next ∼1000 km. DWBC core temperatures increase downstream from ∼0.4°C at the Ammassalik array to ∼1.0°C at our Cape Farewell array (and further to ∼1.4°C off southwest Greenland at ∼61°N). This core warming of 0.6°C between Ammassalik and Cape Farewell is the same as the transport-weighted warming noted above.

If it is assumed that mixing in the DWBC is only associated with entrainment, then the similarity of transport values at the two sites coupled with the downstream warming leads one to conclude that a decrease in transport has occurred between the two measurement periods. If detrainment is present and if fluid near the selected upper boundary (here 3°C) warms and retains downstream momentum, it is necessarily lost to the boundary current as defined. The transport may then decrease with distance downstream and no conclusion can be reached about the comparability of two separated transport estimates.

b. Transport denser than 27.85 kg m−3

Recent transport estimates have been made along the Greenland coast from the sill in Denmark Strait southward. A review of the measurements made in two projects [Variability of Exchanges in the Nordic Seas (VEINS) and Arctic–Subarctic Ocean Fluxes (ASOF)] between 1997 and 2005 may be found in Dickson et al. (2008). Transport estimates were made both at the sill and south of the sill at locations shown in Dickson et al. as Fig. 19.10 and reproduced here as Fig. 13. Sections 24 were the site of the transport arrays and 1–6 of repeated CTD lowerings. Section 6 corresponds closely to the Cape Farewell section that we occupied, and section 3 to the Ammassalik section where Dickson and Brown (1994) made their measurements. Dickson et al. (2008) argue for the essential long-term stability of the overflows, but show that at the sill there was a strengthening in the years from 1997 to 2000 and then a weakening from 2000 to 2002 and subsequent stability. The records end in 2005. However, the picture differs for the transport estimates at sections 24. Table 19.1 of Dickson et al. (2008) shows transports of water denser than 27.85 kg m−3 and the water mass estimates from which they are composed. The transport values have a minimum in 1997 of 2.5 Sv and a maximum in 2005 of 5.0 Sv. Dickson et al. (2008) also mention measurements on the Cape Farewell section (section 6) but exclude them from their summary, asserting that “the transports appear to increase downstream between sections 4 and 6, perhaps due to a recirculation loop within the Irminger Basin.” As described above, we find the transport below a density of 27.85 kg m−3 to have a mean value of 4.5 Sv. While this is within the range (2.5–5.0 Sv) of the upstream values, the argument adduced in the preceding section about mixing and downstream entrainment–detrainment applies here too, and we can draw no definite conclusions about the importance of asynopticity of the estimates.

c. Transport denser than 27.8 kg m−3

We now come to the transport of water below a density of 27.8 kg m−3. Clarke (1984) estimated the transport of dense water south of Cape Farewell (observations made in 1978) and the transport of 13 Sv is generally quoted. To compare our new measurements with his, we attempt to estimate the uncertainty of his value. First, therefore, we review and reevaluate his method and results.

To obtain a transport estimate for the DWBC, Clarke used geostrophic current calculations derived from a CTD section totaling 21 stations between Cape Farewell and Flemish Cap. The geostrophic currents were initially referenced to a level of zero motion at 1500-m water depth. To provide reference current information, he had six current meters on three moorings. Three of the current meters were at depths of 600 m or shallower. The other three, on two of the moorings, were at 1900-m water depth or deeper. The record lengths were ∼60 days. As well as the current meter measurements, additional reference current constraints were obtained from consideration of the likely direction of flow of various water masses. All constraints are summarized in Clarke (1984, Table 3) and his estimated velocity section shown in his Fig. 6.

Having produced his best estimate of the section velocities, he then calculated transports. His surface-to-bottom “Greenland Slope” transport (i.e., the whole boundary current system excluding the shallow shelf waters; see his Table 2) was 33.5 Sv, of which 16.0 Sv made up the DWBC. It is noteworthy that for this original calculation, his Labrador Sea DWBC inflow (off Cape Farewell) and outflow (east of Newfoundland) were in near balance (to within 0.5 Sv). This is what then followed (in brief): (i) noting that the net volume flux across the whole section (i.e., across the Labrador Sea) was unbalanced, (ii) making allowance for unmeasured currents on the Canadian side of the section, and (iii) estimating the expected net imbalance across the section due to inflow to the Labrador Sea from Baffin Bay, he decided that the simplest way to achieve a flux balance was to reduce all reference currents by 10%. This is the cause of the reduction from his best initial estimate of the measured DWBC flux at Cape Farewell of 16 Sv to the now widely accepted value of 13 Sv. The net Greenland Slope transport is also reduced, from 33.5 Sv to 27 Sv. However, as he noted at the time, “such adjustment unbalances the deep- and bottom-water portion of the flow by some 3 × 106 m3 s−1 in the process.” Furthermore, before making this observation, he notes that “it is the upper water masses which are in imbalance.”

What are we to make of this? First, we note that the net boundary current transport off Cape Farewell has been estimated by a number of authors who also included in situ current measurements in their calculations: Bersch (1995) obtained a transport of 32 Sv (including vessel-mounted ADCP data) for a 1991 section; Lherminier et al. (2007) obtained transports of 39 and 32 Sv for sections in 1997 and 2002, including LADCP data; and Holliday et al. (2009) obtain a transport of 40 Sv for a section in 2005 with LADCP data. These transports are consistently between 32 Sv and 40 Sv so Clarke’s 27 Sv (adjusted) looks too low.

Second, we can observe that Clarke’s (unadjusted) 33.5 Sv may also have been too low. A close comparison of his velocity section (his Fig. 6) with a recent analog, Fig. 5 in Holliday et al. (2009), clearly suggests that an artifact of Clarke’s level of zero motion survives in his velocity section at depths between 800- and 1500-m water depth, where isotachs curve into the seabed. The 0, 5, and 10 cm s−1 isotachs in both papers are remarkably similar in location (distance offshore) and (near vertical) orientation; in Holliday et al. (2009), however, the higher-velocity isotachs for 15, 20, and 25 cm s−1 are nearly parallel to those of lower velocities and exhibit no middepth “break.” Taking the modern current profiles to be properly representative of the current structure, Clarke’s resulting underestimate of the boundary current around his break would be roughly 800 m (depth range) × 50 km (width of the relevant portion of the section) × 5–10 cm s−1 (approximate net current bias in the area), a total of 2–4 Sv. His Cape Farewell boundary current flux would now be of order 35.5–37.5 Sv, and his net Labrador Sea flux further out of balance.

Finally, we note that modern measurements of the total boundary current transport off Newfoundland (Fischer et al. 2004) give values of ∼37 Sv, very similar to the Cape Farewell estimates cited above. It is therefore our contention that Clarke’s first DWBC estimate of 16 Sv was his best estimate, and that his belief that the imbalance lay in the upper waters was correct. A uniform reduction in his reference current strengths was the simplest course of action to achieve balance given the paucity of information at the time, but was wrong. Rather, he needed to strengthen his estimates of upper-ocean transports near Newfoundland.

However, the value of 13 Sv has been widely quoted and seemed to have obtained support during the World Ocean Circulation Experiment (WOCE) when Fischer et al. (2004) reported a value of 13.8 Sv for the southward transport of dense water on the southern edge of the Labrador Sea. However, farther south at the western exit of the subpolar gyre, east of the Grand Banks, Schott et al. (2004) measured the DWBC transport from mooring arrays deployed between 1993 and 1995 and between 1999 and 2000. For density >27.8 kg m−3, they obtained a value of 8.9 Sv, essentially equal to that found by us at Cape Farewell.

What then of the difference between our measurement of 9 Sv and Clarke’s 13 or 16 Sv? With ε = 1 Sv, our measurement of 9 Sv lies 4ε from 13 Sv. The probability of the true mean of our measurements being 13 Sv (or more) is vanishingly small.

We also need to address the possible error of Clarke’s value(s). We have seen that the transport decorrelation time scale is ∼10 days, so the sample standard error for a 60-day record (assuming σ = 2.9 Sv) is 0.9 Sv. We have also seen from his calculations that 13 or 16 Sv are defensible values, so we assign a total standard error of 2 Sv (labeled εC) to his estimate. Our value of 9 Sv is then 2εC below Clarke’s 13 Sv (erring on the generous side), so there is a 2.2% chance (assuming Gaussian statistics) that Clarke’s value should have been 9 Sv. To illustrate the chance that our value is biased low while Clarke’s is biased high, the probability that our value should be 11 Sv is 2.2% (the mean is 2ε low), while the probability that Clarke’s value is also 11 Sv is 15.8% (the mean is 1εC high). The joint probability of both occurrences is therefore 0.3%. The probability that the two measurements are not significantly different—that is, cannot be distinguished from one another—is negligible, therefore.

Since we believe that the 1978 and the 2005–06 transports do differ significantly, then what other data exists to illuminate this difference? Between 1978 and 2005–06, the only current meter measurements made near Cape Farewell were made during WOCE. Two moorings with three instruments only were deployed during the period 1995–97 (WOCE current meter archive, National Oceanographic Data Center, and German component of ACM8; information online at http://www.nodc.noaa.gov/). In Fig. 14 we have superimposed these annual-mean vectors at 110 m above bottom on measured and interpolated values at this depth from our 2005–06 results. It will be seen that the WOCE measurements are 20%–30% larger, so we can make a rough estimate of how much extra transport this represents, without altering any other parameters. Assuming they represent a barotropic change in currents of 5–10 cm s−1 over a mean depth range of 1000 m and a horizontal distance of 50 km, then the DWBC transport could have been 2.5–5.0 Sv higher, resulting in a total flux of 11.5–14.0 Sv, potentially placing WOCE-era DWBC transports back in the “accepted” region of 13 Sv. Prior to 2005, numerous geostrophic current estimates have been made in the vicinity of Cape Farewell, and those covering a 40-yr period summarized by Bacon (1998) and Kieke and Rhein (2006). These authors show that transports denser than 27.8 kg m−3 range from 2.5 to 8.5 Sv and are larger by roughly a factor of 2 between the mid-1970s and the mid-1980s than observed before or after. While these reference-level-based transport calculations are imperfect, this is possible confirmation of the weakening of the flow between the two sets of measurements.

d. Variability within 2005–06

Turning now to the time changes seen in our data (in Fig. 12 and Table 4), the repeated WOCE–VEINS–ASOF CTD data on the Cape Farewell section enable us to assess these results. Holfort and Albrecht (2007) captured the variations in the overflow properties there by displaying a time series of the core salinity, defined as the salinity on the density surface σ2 = 37.12 kg m−3. We have reproduced their results post-1995 from their Fig. 5, averaging the data within each cruise and calculating the potential temperature change that goes hand in hand with the salinity change; see our Fig. 15. Apart from emphasizing the temperature, we have also added calculations from Discovery cruises 298 (2005), 309 (2006) and 332 (2008). It will be seen that the core temperatures and salinities from 2006 to 08 are a little higher than the earlier ones, but for the most part the variations are interannual. We believe that by these data we have demonstrated that the current meter temperature changes seen in Table 4 probably do not represent a long-term trend. The transports seen in sections 2–4 of Dickson et al.’s (2008) Fig. 13 and reported in their Table 19.1 also show interannual variability, so we believe that the transport changes we see (also Table 4) within our dataset are probably interannual.

e. A benchmark for numerical and theoretical models

A valuable aspect of the measurements reported here is the benchmark they provide against which numerical and theoretical models can be tested. In the vicinity of the sill, numerical and theoretical models are effective in representing the overflow characteristics and dynamics [see Jungclaus et al. (2008) for a recent review]. However, there are few dedicated models that explore the downstream development of the DWBC. Two such examples are the streamtube model of Price and Baringer (1994), and the reduced-gravity plume model of Jungclaus and Backhaus (1994). Assuming the same overflow at the sill (2.9 Sv), although quite different in structure, they produced the same answer for the enhanced overflow transport in the vicinity of Cape Farewell, namely, 4.1 Sv, close to our observed Denmark Strait overflow DWBC component (density > 27.85 kg m−3) of 4.5 Sv. Clearly, this does not imply that the models are “right”; for example, the Price and Baringer (1994) model allows no mixing farther downstream of the sill than 100 km, whereas Lauderdale et al. (2008) observe significant mixing all along the overflow/DWBC path from the Denmark Strait sill to Cape Farewell and beyond.

On the other hand, for oceanwide (global) numerical models the picture is quite different. Saunders et al. (2008) have recently argued that the level model of Wunsch and Heimbach (2006), even with data assimilation, yields too shallow a meridional overturning in the subtropical gyre of the North Atlantic. The shallowness of the overturning is attributed to the failure of the modeled overflow plume in the subpolar gyre to reach the depths seen here, namely 2000–3000 m. As a further example, the Ocean Circulation and Climate Advanced Modelling Project’s (OCCAM) model (see Lee et al. 2007 for a description), a forward global model of extreme resolution, mixes away the overflow on a “steppy” bottom and a WBC is not found below 800 m at Cape Farewell. On the deep seabed at this location the model flow is very weak and toward the north!

To summarize, our understanding of the northern North Atlantic DWBC system is presently incomplete in terms of forcing, dynamics, and spatial and temporal variability. We look forward to a generation of ocean general circulation models that can throw light on these issues.

Acknowledgments

We wish to acknowledge the support of the master and crew of the RRS Discovery on two cruises (298 and 309) and the technical support of the seagoing engineering group, and in particular Ian Waddington, Dan Comben, and Steve Whittle (moorings and instruments), and Dave Teare (CTD). We are most grateful to IFREMER in the persons of Pascale Lherminier, Nathalie Daniault, and Herlé Mercier for allowing us the use of their data. Our thanks to Liz Kent and Dave Berry for helpful discussions. We are also grateful to two patient reviewers whose comments helped us improve the presentation of our early versions. The Natural Environment Research Council supported the field program under Grant NER/T/S/2002/00453.

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Fig. 1.
Fig. 1.

Map of the locations of the moorings described in this report (Y–G), of two moorings (252 and 253) described in the text (Clarke 1984), and of two moorings to the north (ACM8); see text. The bathymetry is shown in meters.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 2.
Fig. 2.

Cross section of the moorings, showing locations of the instruments. Mean depths of the 3°C isotherm and of the density surfaces σθ = 27.85 and 27.8 kg m−3 are also shown.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 3.
Fig. 3.

(top) Potential temperature (θ) on the mooring line derived from CTD casts made during the deployment of the moorings (4–6 Sep 2005) showing the cold layer on the seabed, and (bottom) the corresponding salinity showing the saline layer above the cold layer. Both are components of the DWBC.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 4.
Fig. 4.

Salinity (upper trace) and potential temperature (lower trace) for the entire record of RCM h4, i.e., the instrument on mooring H at depth 2365 m, 53 m above the seabed. The inset panel shows in detail the first 2.5 days of the salinity record.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 5.
Fig. 5.

Record-mean current vectors at each IFREMER and NOCS mooring: the current increases almost everywhere toward the bottom and closely parallels the bathymetry (depths in m).

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 6.
Fig. 6.

The character of the current in the core of the DWBC (h4, main panel), and on its edge (e3, inset) as seen in virtual displacement plots, both approximately 50 m above the seabed. Elapsed time is marked every 10 days, with every 50 days numbered.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 7.
Fig. 7.

The DWBC speed fluctuations on four instruments (denoted by color) on mooring B in the direction 032°: the means have been removed, and the vertical coherence of the fluctuations measured at four locations over the depth range 1980–2600 m is striking.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 8.
Fig. 8.

Mean value of the transport per unit width, on each mooring, colder than 3°C (long dash) and denser than 27.85 kg m−3 (short dash), with estimates of the likely errors of the former.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 9.
Fig. 9.

Time series lasting 284 days of the 12-h average DWBC total transport colder than 3°C. The superimposed circles are 20-day-average values.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 10.
Fig. 10.

Spectral analysis of the total transport colder than 3°C displayed in variance-preserving manner and similar to figures shown for transports at upstream locations (Saunders 2001). The logarithm of the frequency (cpd) (logfreq, primary x axis) is plotted, and the equivalent period (days) is shown on the secondary x axis.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 11.
Fig. 11.

Contributions to the DWBC transport (Sv) colder than 3°C from each mooring for (a) the record average and for extreme conditions: two cases of (b),(c) high and (d),(e) low transport. Date range and total transport are shown for each panel.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 12.
Fig. 12.

Potential temperature vs salinity for CTD stations on the mooring line: data from the NOCS deployment cruise (298, 4–6 Sep 2005, black) and the NOCS recovery cruise (309, 22–31 Aug 2006, red). The warming and salinity increase below 3°C is evident. The three density lines are described in the text.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 13.
Fig. 13.

CTD station positions for VEINS cruises. The data shown are for the 1997 cruise and were obtained from the NODC WOCE archive.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 14.
Fig. 14.

Year-long current vectors on the Cape Farewell section approximately 110 m above the seabed. The black vectors refer to the data reported here (moorings Y–C), and the gray vectors to data from 1995 and 1996. See text for source (WOCE).

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Fig. 15.
Fig. 15.

Core properties in the overflow on the density surface σ2 = 37.12 kg m−3 derived from the NODC WOCE data archive, the data of Holfort and Albrecht (2007), and after 2004 recent Discovery cruises.

Citation: Journal of Physical Oceanography 40, 4; 10.1175/2009JPO4091.1

Table 1.

Mooring details, showing position, water depth, and instrument depth. The sample intervals are 30 min for ACM 11 and 60 min for ACMs 7 and 8. The moorings designated Y and Z are IFREMER moorings; the remainder are NOCS moorings.

Table 1.
Table 2.

Statistics of the record-mean low-pass currents and temperatures. The means commence 6 Sep 2005 and end 17–21 Jun 2006 for moorings Y and Z data (the property of IFREMER), while for the remainder they end 23–30 Aug 2006. Instruments h2 and b4 were so labeled before the discovery that they yielded no data and the z5 current direction was adjusted; see text. Ident is the instrument identifier, θ the potential temperature (°C), V032 the speed normal to the mooring line, and U122 the speed parallel to the mooring line. The mean current vector is calculated from the two components and its magnitude is the speed; Dir is the direction (°true, clockwise relative to north). The final column shows the number of 12-h records. Rms is the standard deviation of the relevant variable.

Table 2.
Table 3.

Details of the data employed in calculating transports. For each of three cases [density (σθ) greater than either 27.8 or 27.85 kg m−3, and potential temperature (θ) less than 3°C], the height (Ht) above seabed (m) for the relevant surface on each mooring, the transport per unit length (Trans. PUL, in units of m2 s−1), and the net transport (Trans. net, Sv) are shown. The mooring identifier (mrng. ident.) is shown in the first column, and the second column shows the effective width (eff. width, km) assigned to the mooring. For moorings Y and Z the values are 9½-month means; for the remainder, 11½-month means.

Table 3.
Table 4.

Warming in temperature and weakening in transport colder than 3°C. The first 60 days started on 6 Sep 2005 and the last ended on 18 Jun 2006. Transport (Sv) is in the direction 032°T.

Table 4.
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