Stationary Sea Surface Height Anomalies in Cyclonic Boundary Currents

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Stationary Sea Surface Height Anomalies in Cyclonic Boundary Currents: Conservation of Potential Vorticity and Deviations from Strict Topographic Steering SARA BROOMÉ AND JOHAN NILSSON Stockholm University, Stockholm, Sweden (Manuscript received 11 November 2015, in final form 14 April 2016) ABSTRACT In high-latitude subpolar seas, such as the Nordic seas and the Labrador Sea, time-mean geostrophic currents mediate the bulk of the meridional oceanic heat transport. These currents are primarily encountered along the continental slopes as intense cyclonic boundary currents, which, because of the relatively weak stratification, should be strongly steered by the bottom topography. However, analyses of hydrographic and satellite altimetric data along depth contours in Nordic seas boundary currents reveal some remarkable, stationary, along-stream variations in the depth-integrated buoyancy and bottom pressure. A closer examination shows that these variations are linked to changes in steepness and curvature of the continental slope. To examine the underlying dynamics, a steady-state model of a cyclonic stratified boundary current over a topographic slope is developed in the limit of small Rossby numbers. Based on potential vorticity conservation, equations for the zeroth- and firstorder pressure and buoyancy fields are derived. To the lowest order, the flow is completely aligned with the bottom topography. However, the first-order results show that where the lowest-order flow increases (decreases) its relative vorticity along a depth contour, the first-order pressure and depthintegrated buoyancy increase (decrease). This response is associated with cross-isobath flows, which induce stretching/compression of fluid elements that compensates for the changes in relative vorticity. The model-predicted along-isobath variations in pressure and depth-integrated buoyancy are comparable in magnitude to the ones found in the observational data from the Nordics Seas.

1. Introduction The continents of the Northern Hemisphere provide boundaries for the oceans that allow the oceanic meridional heat transport to largely be brought about by time-mean geostrophic currents. This can be compared to the Southern Ocean where the lack of boundaries yields a situation where meridional heat transport instead is mainly mediated by eddies (Döös and Webb 1994; Marshall and Speer 2012). Furthermore, the relatively weak stratification of northern high-latitude

Denotes Open Access content.

Corresponding author address: Sara Broomé, Department of Meteorology, Stockholm University, SE-106 91 Stockholm, Sweden. E-mail: [email protected] DOI: 10.1175/JPO-D-15-0219.1 Ó 2016 American Meteorological Society

ocean basins causes the geostrophic currents to follow the underlying bottom topography. Such is the case in the Nordic seas and the Labrador Sea, where the largescale circulation is dominated by time-mean boundary currents, largely tied to the continental slopes. These basins connect the North Atlantic with the Arctic Ocean; the warm and salty Atlantic water that flows along the eastern boundary of the Nordic seas is the main source water for the Arctic Ocean and thus affects the Arctic climate and sea ice (e.g., Zhang et al. 1998; Orvik and Niiler 2002; Carton et al. 2011). These subpolar seas are also regions of deep-water formation, feeding the lower limb of the Atlantic meridional overturning circulation (AMOC). That the boundary currents play an important role in the production of deep overflow waters from the Nordic seas that form the North Atlantic Deep Water has been suggested by, for example, Mauritzen (1996), Eldevik et al. (2009), and Isachsen et al. (2007). Similarly, Böning et al. (1996) and

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FIG. 1. Map of the Nordic seas with ocean depth (m). Indicated are the Lofoten Basin (LB), Vøring Plateau (VP), Barents Sea opening (BSO), location of the Svinøy hydrographic section, and parts of the Greenland–Scotland Ridge. In red are also indicated rough pathways of the NwAFC and the NwASC and its fractionation at the BSO.

Spall and Pickart (2001) found in modeling studies that the sinking and formation of the middepth Labrador Sea Water takes place near lateral topographic boundaries. In the present work we focus on the Nordic seas, situated in the northeastern corner of the North Atlantic between Greenland and Norway (see Fig. 1). The eastern part of the Nordic seas is mainly influenced by warm and saline Atlantic water, while the western part is filled with cold and freshwater of mainly Arctic origin (Mauritzen 1996). There is a clear separation between the two halves along the topographic ridge running south–north through the seas. The Atlantic water enters the Nordic seas over the Greenland–Scotland Ridge and then flows northward in an eastern boundary current, connecting the North Atlantic Ocean with the Arctic Ocean. The current can be seen to follow the contours of the bottom topography (Helland-Hansen and Nansen 1909; Poulain et al. 1996; Jakobsen et al. 2003; Søiland et al. 2008). There are two main branches of the current: one inner, barotropic branch strictly following the Norwegian continental slope [Norwegian Atlantic Slope Current (NwASC)] and one outer frontal branch [Norwegian Atlantic Front Current (NwAFC)]. The bulk of the transport is in the NwASC; Mork and Skagseth (2010) estimate the time-mean volume fluxes in the Svinøy

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section to be 3.4 and 1.4 Sv (1 Sv [ 106 m3 s21), respectively, for the two branches. For a more detailed description of the circulation in the Nordic seas, see, for example, Mork and Skagseth (2010), Orvik and Niiler (2002) and Jakobsen et al. (2003). In contrast to the stratified, low-latitude ocean, the Sverdrup balance (Sverdrup 1947) does not generally provide a leading-order description of the circulation in the weakly stratified subpolar oceans. Here, the gradients of planetary potential vorticity f/H, where f is the Coriolis parameter, are strongly dominated by the steep topography H, causing the flow to trace the isobaths around the basins (e.g., Orvik and Niiler 2002; Jakobsen et al. 2003; Koszalka et al. 2011). The underlying dynamics reflects the conservation of potential vorticity of the low Rossby number flows for which the leadingorder bottom velocity is directed along the f/H contours (Greenspan 1968). Nøst and Isachsen (2003) showed that the leading-order bottom flow along closed f/H contours are primarily set by a global vorticity balance in which the along-contour, integrated surface wind stress and bottom stress balance each other. The integrated wind stress over the closed f/H contours in the Nordic seas and the Arctic Ocean tends to be cyclonic (i.e., dominated by an Iceland low pressure), yielding a prevalence of cyclonic, along-isobath bottom currents in the basins (Nøst and Isachsen 2003; Rudels et al. 2000; Aaboe and Nøst 2008). By cyclonic, we refer to flows with shallow water to the right, sometimes also referred to as prograde (see, e.g., Li and McClimans 2000). On the open f/H contours, which are typically shallower than about 500 m, the bottom velocities on the slope in the Atlantic can be viewed as upstream boundary conditions for the Nordic seas. The downstream changes of the bottom velocities on these f/H contours are mainly related to bottom buoyancy variations (Walin et al. 2004; Schlichtholz 2007; Spall 2005; Aaboe and Nøst 2008); if the buoyancy decreases downstream, the bottom velocity becomes increasingly cyclonic, making the flow more barotropic (Walin et al. 2004). There is, however, a profound asymmetry in the tendency to follow f/H contours between cyclonic (prograde) and anticyclonic (retrograde) flows (e.g., Nøst et al. 2008). One reason is that an anticyclonic current along a continental slope, which flows with shallow water to the left, supports stationary topographic waves. As a result, along-stream bathymetric variations tend to generate stationary waves, which can create significant flow across the depth contours. In fact, most well-known examples of boundary currents that separate, such as the Gulf Stream, are anticyclonic flows (see, e.g., Marshall and Tansley 2001). We note, however, that Björk et al. (2010) report hydrographic data that indicate partial

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separation of a cyclonic boundary current as it passes over some strong topographic gradients on the Morris Jesup Rise off North Greenland. As the time-mean flow in the Nordic seas is cyclonic and, to a good approximation, in geostrophic balance, the theory summarized above suggests that the timemean bottom pressure should essentially be aligned with the topography. In the present work, we analyze the time-mean circulation in the Nordic seas on the basis of hydrographic and altimetric data, with a particular focus on the NwASC [see, e.g., Skagseth et al. (2004) for altimetry in the Svinøy section]. In the data analysis, we partition the dynamic surface pressure into a bottom pressure contribution and a steric height contribution, where the latter is proportional to the depth-integrated buoyancy. Although the main variations of these dynamical fields occur perpendicularly to the isobaths, there are some significant variations along isobaths. For instance, along the isobaths beneath the NwASC the variations in the altimetrically derived surface height suggest cross-isobath volume transport on the order of a few Sverdrups, which flow toward deeper water in some regions and back shoreward in other regions. This could be reflecting errors in the data, but we argue that these along-isobath pressure variations can be explained, at least qualitatively, by the conservation of potential vorticity of a low Rossby number flow that experiences along-stream changes in the steepness and curvature of the topography. This result provides, for example, information on the timemean exchange of tracers across the depth contours in cyclonic boundary currents, for example, the exchange between the two branches of the Norwegian Atlantic Current (see Raj et al. 2015). The remainder of the paper is organized as follows: Section 2 describes the decomposition of the geostrophic velocity into a thermal wind and a bottom velocity component and also presents the satellite altimetry and hydrographic data. In section 3, the results of the observational analysis are presented. In section 4, a steady-state, potential vorticity–conserving model of flow over along-stream varying topography is presented, and in the appendix this is complemented with an analysis of a linearized quasigeostrophic model. In section 5, the model is discussed in relation to the observational results. Some broader implications are discussed in the concluding section 6.

2. Method and data a. Thermal wind and bottom velocity We consider a flow in hydrostatic balance with the density field given by

2439 r 5 r0 1 r0 ,

(1)

where r0 is a reference density, and r0 is an anomaly. We further define the buoyancy anomaly as b [ 2g

r0 , r0

(2)

where g is the gravitational constant. The hydrostatic equation is then given by ›f 5 b, ›z

(3)

where f is the dynamic pressure. Integrating the hydrostatic equation gives f(x, y, z, t) 5 gh 2

ð0 b dz,

(4)

z

where we have used the boundary condition that the pressure at the surface is given by the dynamic surface height h. As defined here, h corresponds to the absolute dynamic topography data, as described below, that among several corrections includes the atmospheric pressure anomaly. Equation (4) represents the most common way to write the hydrostatic pressure; there is a surface height added to the weight related to the integrated buoyancy anomaly above the level z. However, it can be dynamically revealing to decompose the dynamic pressure into one bottom pressure component and one component related to the buoyancy field. Fofonoff (1962) and several subsequent researchers have used such a decomposition of the pressure field. Following Fofonoff, we define a dynamic bottom pressure as fB (x, y, t) [ f(x, y, z 5 2H, t) 5 gh 2

ð0 b dz,

(5)

2H

where H is the depth. Using this, we can write the dynamic pressure in Eq. (4) as ðz f(x, y, z, t) 5

2H

b dz 1 fB ,

(6)

that is, the pressure at level z is given by subtracting the contribution from the buoyancy in the layer between z 5 2H and z from the bottom pressure. The introduction of the dynamic bottom pressure fB in Eq. (5) allows the free-surface height to be partitioned as follows: h 5 hS 1 hB , where the steric height contribution hS is defined as

(7)

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hS [

1 g

ð0 b dz,

(8)

2H

and the contribution related to the dynamic bottom pressure is 1 hB [ fB . g

(9)

In the data analysis, we assume flows that are in geostrophic balance, that is, u(x, y, z, t) 5 f 21 k 3 =f ,

(10)

where u is the horizontal velocity and k is the vertical unit vector. We here comment on how the dynamic bottom pressure can be used to infer cross-isobath flow. By using Eq. (6), we obtain u(x, y, z, t) 5 f 21 k 3 =

ðz 2H

b dz 1 f 21 k 3 =fB .

(11)

By using Leibnitz’s rule for differentiation of integrals, we can rewrite the geostrophic velocity as u(x, y, z, t) 5 f

21

2H

Here, we do not explicitly consider the along-isobath bottom velocities, and in section 5 we consider only fB that governs the cross-isobath flow.

b. Absolute dynamic topography and bottom topography data The observational portion of this study is based on a satellite altimeter product produced by Ssalto/Duacs and distributed by AVISO, with support from CNES (http://www.aviso.altimetry.fr/duacs/). The product is called absolute dynamic topography (ADT) and depicts the dynamic sea surface height, that is, the sea surface height due to motion; a geostrophic approach on the ADT gives the surface geostrophic current. We use the global gridded data on a 1/ 48 Cartesian grid provided as daily fields from January 1993 to May 2014. All data presented here are the climatologic time mean over all available data. For some calculations and figures, we use global bathymetry and elevation data (Becker et al. 2009) at a resolution of 0.058 3 0.058. The ADT is here used as an observation of h in Eq. (7). Accordingly, the ADT could be divided into a steric and a bottom pressure contribution. To this end, a steric height is constructed from hydrographic data.

c. Hydrography, steric height, and bottom pressure

ðz k3

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=b dz 1 uB (x, y, t).

(12)

Here, we have introduced the bottom velocity component uB, which is defined as uB (x, y, t) [ f 21 bB k 3 =H 1 f 21 k 3 =fB ,

(13)

where bB [ b(x, y, z 5 2H, t) is the buoyancy at the bottom. In Eq. (12), the first term is the thermal wind velocity relative to the bottom, which is referred to as the ‘‘baroclinic’’ flow component by Fofonoff (1962). The bottom velocity is composed of two terms. The second term is the velocity associated with the bottom pressure gradient. This is, however, not a gradient at fixed z; the first term compensates for this, which can be seen by noting that at the bottom =zf 5 =fB 1 (›f/›z)=H, where =z is the gradient at fixed z and that from the hydrostatic relation (›f/›z) 5 bB. The compensating term is directed along the depth contours and flows anticyclonically (with shallow water to the left) if bB is positive. Mathematically, Eq. (13) is well defined, but in databased calculations care has to be taken as the bottom velocity can be the difference between two potentially large terms [see Aaboe and Nøst (2008) for a computation of bottom velocities]. However, we note that cross-isobath velocities still rely on a gradient at fixed z.

The hydrographic data are the World Ocean Atlas 2013 (WOA13) climatology of temperature (Locarnini et al. 2013) and salinity (Zweng et al. 2013) on a 1/ 48 Cartesian grid. For comparison, an alternative hydrography is also used and shown in selected figures. This dataset was created using isopycnal mapping by Isachsen and Nøst (2012) and covers the Nordic seas at 50–70-km horizontal resolution. Over a flat sea bottom, the reference density used to define the steric height is dynamically irrelevant. When the bottom depth H varies, however, the partitioning of the surface pressure gradient into a steric and a bottom component is no longer unique; lateral gradients of steric height over variable bottom depths gives a contribution proportional to the bottom slope that depends on the choice of the reference buoyancy. For example, by changing the reference buoyancy according to b / b 1 bref, where bref is a constant, it follows from Eq. (8) that the steric height and the bottom pressure become hS / hS 1 brefH/g and hB / hB 1 brefH/g, respectively. By computing the buoyancy relative to a wisely chosen z-dependent reference density r0(z), we can still obtain dynamically relevant steric heights and bottom pressure fields. A sensible requirement is that when the gradient of h is zero, that is, no surface geostrophic flow, the gradients of hS and hB should vanish independently, or at least be small, since two large terms of opposite signs

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is southward and strong near the coast and alternates direction over the slope. For depths greater than 1000 m, the two methods yield very similar steric height gradients. Thus, we argue that our specification of r0(z) provides a reasonable definition of the steric height over sloping bottoms that contains relevant dynamical information. Having defined the reference density, the contribution to the free-surface height from the bottom pressure is now constructed as the difference between the timemean ADT and steric height: hB 5 hADT 2 hS .

FIG. 2. The steric height (m) computed with two different reference densities: one constant potential density r0(S0, u0), where S0 and u0 represent deep-water properties (hatched), and one z-dependent reference density (solid). The steric heights are shown in a section across the slope along 668N, where the z-dependent reference density is defined. In black is the section’s bathymetry.

would be cumbersome. In a single x–z section, we can achieve this by selecting the reference density profile as r0 (z) 5 rB (x, z 5 2H);

(14)

thus, the vertical reference density profile equals the bottom density on the slope at each depth. This choice gives hS 5 0 if the isopycnals are horizontal in the section. Furthermore, by definition, the bottom buoyancy bB will be zero in the section, and there will be no cancelling terms in the bottom velocity equation [Eq. (13)]. Here, we choose r0(z) to be defined in a section at 668N, just south of the Vøring Plateau, which has a fairly gentle slope and thus allows for a better resolution of the bottom buoyancy in the gridded data compared with sections having steeper slopes. The bottom value of the density is taken as the lowest available level in the hydrographic data for every location in the section. The bottom density varies slightly upstream and downstream of the 668N section, implying some bottom buoyancy variations that affect the along-slope bottom velocity [Eq. (13)] [see Aaboe and Nøst (2008) for further discussion]. Figure 2 shows the steric height in the 668N section computed using both the reference density profile r0(z), described by Eq. (14), and a constant reference density. The steric height based on r0(z) increases monotonically toward the coast and, accordingly, tends to capture the thermal wind relative to the bottom also in shallower waters. The steric height based on a constant reference density, on the other hand, implies thermal wind flow that

(15)

Finally, we emphasize that the variations in steric height and bottom pressure along a depth contour, to be discussed in section 5, are independent of the choice of reference density—it matters only for variations across the depth contours.

3. Observational analysis Here, we present the decomposition of the satellite data of ADT into contributions from steric height and bottom pressure, as discussed theoretically in section 2a. To begin with, we describe briefly some qualitative features of these fields in the Nordic seas and then go on to a more quantitative discussion of the pressure fields along the NwASC. It is worth noting that Johannessen et al. (2014) made a similar decomposition of the surface height field into a steric and a bottom component. Differences lie in the choice of hydrographic and ADT data (where the ADT was smoothed using an 80-km Gaussian filter) and in how they define the steric height; they use a constant reference depth of 1500 m and present a steric height field where areas shallower than 1500 m are filled in with steric heights with reference depths every 100- to 500-m depth. Our approach instead uses a varying depth but lets the reference buoyancy vary with vertical coordinate. The resulting steric height in Johannessen et al. (2014) is broadly similar to, but smoother than, the one presented here. Figure 3 shows the time-mean satellite altimetric ADT, the hydrographic steric height, and the resulting bottom pressure in the Nordic seas. The time-mean ADT and its associated geostrophic velocities have been discussed by, for instance, Raj (2013) and Chafik (2014). They, and Chafik et al. (2015), also discuss the variability and the ADT-based eddy kinetic energies, where the highest activity is found in the Lofoten Basin (LB) and on the slope east of it. The mean ADT field roughly outlines a cyclonic surface circulation that tends to be aligned with the topography, and the steepest gradient

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FIG. 3. ADT and its two contributing components’ steric height and bottom pressure (m). Note that the color scale of the steric height and the bottom pressure has been offset by a dynamically irrelevant constant to have the same range in all panels. Outlined are the contours of bottom depths of 500, 750, 1000, 1500, 2000, and 3000 m.

can be found along the Norwegian continental slope, signifying the geostrophic velocity of the NwASC [cf. Fig. 1 and Mork and Skagseth (2010)]. The steric height, which is proportional to the vertically integrated buoyancy anomaly, has in general stronger along-isobath variations than the ADT. Specifically notable is the maximum in steric height in the Lofoten Basin and the associated contours crossing the isobaths of the continental slope. While a geostrophic approach on the ADT and the steric height gives the surface and thermal wind velocities, respectively, it should be noted that one needs to be careful in interpreting velocities from the bottom pressure in Fig. 3; as in Eq. (13), a lateral gradient would not directly be a gradient at fixed z. However, cross-slope velocities associated with gradients on constant depth (as analyzed in section 5) are still well defined. From theoretical considerations and data-based simple modeling (Nøst and Isachsen 2003; Aaboe and Nøst 2008), one expects the bottom pressure field in the Nordic seas to be essentially aligned with the topography and associated with cyclonic circulation. This notion is broadly consistent with the bottom pressure shown in Fig. 3. Specifically, the NwASC is traced out by a gradient in the bottom pressure farther northward than in the ADT. To make a more quantitative analysis of the alignment, we estimate the angle between the gradient of the topography and the gradient of the bottom pressure, ADT, and steric height. The alignment is evaluated along depth contours, tracing the NwASC, by constructing

distribution functions (not shown) where complete alignment is represented by a zero. The distributions for the bottom pressure and the ADT are found to be centered around zero, while the steric height distribution is shifted and broader. The peak of the distribution for the bottom pressure is generally found to be higher than that of the ADT, indicating that it is somewhat better aligned with the topography. However, there are some other notable, and more surprising, features in the bottom pressure. To begin with, the bottom pressure indicates a region of anticyclonic (i.e., southward) flow around the 1500-m isobath that extends from around 628N, 28W and northward between the two current branches toward the Vøring Plateau (VP). A coincident but weaker southward flow is also present in the ADT. Mork and Skagseth (2010) report weak southward, near-surface velocities in this area from satellite and ADCP data, and the near-bottom current meter data described by Aaboe and Nøst (2008) also include a few measurements of anticyclonic flows (i.e., southwestward flow) in this region around 638N. Second, in the deep LB, the bottom pressure suggests an anticyclonic flow at the bottom, collocated with the surface anticyclonic flow seen in the ADT. There resides a semipermanent eddy in the LB that has been discussed before; see, for example, Søiland and Rossby (2013) and Köhl (2007) on the Lofoten Basin eddy and also Orvik (2004) on the deepening of the Atlantic water in the LB. The bottom pressure calculated by Johannessen et al. (2014) suggests, however, a weak

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cyclonic flow at the bottom in the central LB. Also, Nøst and Isachsen (2003) predicted a cyclonic circulation in the LB with their linear model driven by wind stress. If the bottom circulation is indeed anticyclonic, and hence dissipates anticyclonic vorticity in the bottom boundary layer, then there must be an eddy flux of anticyclonic vorticity into the central LB that is stronger than the time-mean cyclonic vorticity input due to the surface wind stress (Nøst and Isachsen 2003). Although we mention these interesting features, the limitations of the data and the present approach require further work to assess their significance. Thus, we stress that these regional features should be interpreted cautiously and do not pursue them further here.

Along-slope steric height variations As mentioned above, the ADT, and in particular the steric height, indicate cross-isobath transport along sections of the NwASC. Also, Chafik et al. (2015) found along-stream variations in their ADT-based volume transport estimate of the NwASC in the 500- to 900-m depth interval. To examine this further, we construct an along-isobath measure of the pressure variations, where the data are interpolated to depth contours of 500, 700, and 900 m along the NwASC. The results are shown in Fig. 4, revealing two local maxima along the slope. These are encountered near the Svinøy section, at about 648N, and off the Lofoten Islands at about 708N. The along-slope variations are found in both the steric height and the bottom pressure, although less accentuated in the latter. This suggests that the somewhat weaker along-isobath variations of the bottom pressure are amplified toward the surface by the baroclinic pressure field, resulting in a larger response in the surface pressure. This corresponds well to the notion of stronger topographic control of the bottom velocity and that the steric height, or the thermal wind, is accountable for most of the cross-isobath transport. Figure 4 also shows the steric height and the bottom pressure calculated from the alternative isopycnal hydrography, revealing broadly similar along-slope variations. From the perspective of an essentially isobathfollowing current, the downstream local maxima in depth-integrated buoyancy along the NwASC, suggested in Fig. 4, are somewhat unexpected. Stationary topographic waves are not present in cyclonic boundary currents and thus cannot account for the variations, which must have a different cause. Possible reasons for the variations in depth-integrated buoyancy are 1) shortcomings in the analysis and erroneous data, for example, errors from uncertainties in the geoid (Johannessen et al. 2014; Rio et al. 2011); and

FIG. 4. ADT, steric height, and bottom pressure (m) along the slope, smoothed by a 50-pt running mean, approximately following the NwASC in Fig. 1 from 628 to 728N. Shown are the deviations from the mean along the slope. ‘‘Along the slope’’ here means that the data have been linearly interpolated to the 500-, 700-, and 900-m depth contours, after which a mean across these contours has been taken, representing a mean across the slope current. The black lines are the WOA13 hydrography, and the gray lines are the isopycnal hydrography.

2) the time-mean flow essentially follows the isobaths, but eddy-induced buoyancy fluxes and/or surface buoyancy fluxes cause the along-stream buoyancy to vary.

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We will return to these possibilities in section 6. In the next section, however, we will show that the variations in bottom pressure and steric height along the isobaths can be explained qualitatively by a model of an essentially adiabatic and isobath-following current that conserves its potential vorticity.

We examine if an inviscid adiabatic model can explain some aspects of the along-stream surface height variations that are suggested by the data. We consider a cyclonic boundary current where the flow is characterized by a small Rossby number and essentially traces the topography. However, conservation of potential vorticity generally requires some cross-isobath velocities where the slope and curvature of the topography changes in the along-stream direction. For a cyclonic boundary current, where there are no stationary topographic waves, the resulting dynamics can be investigated using a local perturbation analysis as will be illustrated below. In contrast, for an anticyclonic boundary current that has the coast to the left, there exist nonlocal, free, stationary topographic waves, which makes a local perturbation analysis unfeasible (see appendix). We formulate a steady model that neglects fluxes from time-dependent eddies in the momentum and buoyancy budgets. Further, we take the Ekman number to be small compared to the Rossby number, which is a reasonable approximation over steep topography (e.g., Nøst and Isachsen 2003; Nøst et al. 2008; and references therein). Specifically, we consider the f-plane dynamics of a steady inviscid flow in hydrostatic balance, governed by the momentum equation ›u ( f 1 z)k 3 u 1 w 5 2=(f 1 juj2 /2) , ›z

(16)

where z [ ›y/›x 2 ›u/›y is the vertical component of the relative vorticity, and w is the vertical velocity. The pressure obeys the hydrostatic relation [Eq. (3)], and the buoyancy equation is ›b 5 0, ›z

(17)

where b here includes the full vertical variation of the buoyancy field. The continuity equation is given by = u1

where we assume a rigid lid. The above set of equations conserves the Ertel potential vorticity Q along streamlines (Allen 1991; Vallis 2006): v =Q 5 0,

(20)

where v [ (u, y, w) and

4. A model along isobaths

u =b 1 w

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›w 5 0, ›z

(18)

Q [ (kf 1 = 3 u) =b .

(21)

We will here use a version of the equation for the vertical component of the vorticity, which is given by k =3[Eq. (16)]: u =(z 1 f ) 5 (z 1 f )

›w , ›z

(22)

where we have omitted the terms related to the curl of w›u/›z as they are negligible in the low Rossby number regime studied below. We introduce a Rossby number Ro [ ^z/f , where ^z is a characteristic value of the relative vorticity, and we take ^ 5 u^(H/L), where u^ is the scale of the vertical velocity as w the horizontal velocity scale and H/L is the aspect ratio. Assuming that Ro 1, we make a series expansion of Eq. (16) in this quantity (see, e.g., Allen 1991; Vallis 2006). To the zeroth order, we obtain in dimensional form u0 5 f 21 k 3 =f0 .

(23)

As the zeroth-order horizontal flow is divergence free, that is, = u0 5 0, the continuity equation [Eq. (18)] reduces to ›w0 5 0. ›z

(24)

Thus, the rigid-lid boundary condition [Eq. (19)] implies that w0 5 0. The bottom boundary condition does not enter until the first order. To the first order, we obtain u1 5 f 21 k 3 =B1 2 u0 z0 /f ,

(25)

where we have introduced the Bernoulli function B1 [ f1 1

ju0 j2 , 2

(26)

which is composed of the first-order pressure and the zeroth-order kinetic energy. The first-order vorticity [Eq. (22)] is

subject to the boundary conditions w(z 5 0) 5 0, w(z 5 2H) 5 2u(z 5 2H) =H ,

(19)

u0 =z0 5 f

›w1 . ›z

(27)

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In addition to the Rossby number, another important parameter for the dynamics is the bottom slope. We define a nondimensional slope parameter ^, d [ DH/H

(28)

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with u 0 according to Eq. (23). This gives (taking f constant) f0 5 f0 (H) ,

(35)

^ is a typical depth. where DH is the depth variation and H The bottom boundary condition [Eq. (19)], in nondimensional form, then becomes

that is, to the lowest order the flow follows the topography, and the vertical velocity is zero. To the first order in the Rossby number, conservation of potential vorticity is given by

Row1 5 2du0 =H 2 Rodu1 =H .

z f 1 u0 = 0 5 0 , u1 = H H

(29)

Thus, if d is of the same order as the Rossby number, implying a weak slope, then the bottom boundary condition [Eq. (19)] becomes to the first order w1 (z 5 2H) 5 2u0 (z 5 2H) =H .

(30)

This is the quasigeostrophic limit where the lowestorder bottom geostrophic velocity is related to w1. However, the continental slopes generally have order one depth variations, implying that d ; 1. In this case, Eq. (29) gives u0 (z 5 2H) =H 5 0,

(31)

implying that the leading-order bottom velocities are aligned with the topography, and w1 (z 5 2H) 5 2u1 (z 5 2H) =H .

(32)

Thus, the ageostrophic velocity enters in the boundary condition at w1. We will here examine this regime. However, for stratified flows, we will show that the quasigeostrophic equations yield similar results if the slope parameter d is larger than the Rossby number but still smaller than unity, an issue that is further discussed in the appendix.

a. The homogeneous limit It is instructive to first consider the homogeneous shallow-water case, which will be useful for interpreting the general stratified case later. In this case, the barotropic potential vorticity is conserved along the streamlines according to the shallow-water version of Eq. (22): z1f 5 0. (33) u= H To the lowest order in the Rossby number, conservation of potential vorticity reduces to conservation of the planetary potential vorticity f/H: f 5 0, (34) u0 = H

(36)

or, since f is constant and u0 =H 5 0, equivalently u1 =H 2 Hu0 =

z0 5 0. f

(37)

Physically, Eq. (37) states that the first-order velocity will be directed toward deeper (shallower) depths— corresponding to lower (higher) values of the zerothorder potential vorticity f/H—when the zeroth-order relative vorticity increases (decreases) downstream along the isobaths. This iterative approximation of potential vorticity conservation is illustrated in Fig. 5; to the leading order, the flow follows strictly the topography as required by Eq. (34). However, where the slope steepens or the isobaths curve, anomalies in relative vorticity arise along the isobaths (Fig. 5a). Equation (37) relates these relative vorticity anomalies to first-order, cross-isobath velocities, which induce compensating changes in the planetary vorticity f/H. This yields conservation of potential vorticity to the first order [see Eq. (36)]. The sum of the zeroth- and first-order flow yield the transport streamfunction depicted by the gray lines in Fig. 5b. As discussed by Nøst et al. (2008), the conservation of potential vorticity implies that in low Rossby number cyclonic boundary currents, the curvature of the streamlines tend to be smaller than that of the topography. Thus, the streamlines tend to follow a somewhat low-passed filtered topography [see LaCasce et al. (2008) and the appendix]. We can obtain an equation for how B1 [and f1 through Eq. (26)] vary along the depth contours. We start by inserting Eq. (35) into Eq. (23), giving that 1 u0 5 (df0 /dH)k 3 =H . f

(38)

The properties of the scalar triple product, stating that a (b 3 c) 5 b (c 3 a) 5 c (a 3 b), can be utilized so that u1 =H can be written with all terms proportional to k 3 =H. Inserting Eqs. (25) and (38) into Eq. (37) then yields

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VOLUME 46

FIG. 5. Theoretically calculated (a) vorticity anomaly and (b) streamfunction for a cyclonic flow with df0/dH constant at 1.6 3 103 m s2 along an idealized slope of the form H(x, y) 5 Hmaxf1 2 exp[2c(x/L)/( y/L)2]g, where c(x/L) is a function specifying the width of the slope. The shading in (a) is the along-isobath vorticity anomaly relative to the upstream value at x 5 0. The black contours in (a) are the isobaths (0.1, 0.3, 0.5, 0.7, and 0.9) Hmax, and the gray line illustrates the upstream velocity profile at x 5 0, which is proportional to the slope j=Hj [see Eq. (38)]. The black contours in (b) are the zeroth-order streamlines, which coincide with the depth contours. The gray contours are the streamlines resulting from the first-order correction, obtained from Eq. (41). These corrected streamlines compensate for the vorticity anomaly in (a) to ensure conservation of potential vorticity.

z0 H df0 t = B1 1 5 0, f dH

(39)

where we have introduced the vector t [ k 3 =H ,

(40)

which is parallel to the isobaths. The quantity within the square brackets is invariant along the zeroth-order streamlines, that is, along constant H. For a cyclonic flow, having shallower water to the right (f0 decreasing offshore), implying that df0/dH , 0, B1 will be higher where z0 is anomalously positive. The transport streamfunction C1, describing the vertically integrated, first-order flow, can be obtained from Eq. (25) by integrating it vertically and multiplying with =H. Some manipulations then yield ð 1 0 B1 dz 5 0. t = C1 2 f 2H

Thus, the streamfunction is proportional to the vertically integrated Bernoulli function; this is a result that also applies for baroclinic flows if u0 is parallel to the isobaths at all depths. The zeroth-order vorticity and kinetic energy can be found by taking, respectively, the curl and square of Eq. (38), resulting in ›f0 2 ›2 f0 (=H)2 , = H 1 f 21 ›H ›H 2 ju0 j2 (=H)2 ›f0 2 5 . 2 2f 2 ›H z0 5 f 21

(42)

This shows that the variations of z0 and kinetic energy along the depth contours are related to variations in the curvature =2H and slope j=Hj of the bottom topography.

b. The stratified case (41)

As in the homogeneous case, we consider flows that are strongly steered by the topography. However, even

BROOMÉ AND NILSSON

AUGUST 2016

in the limit of a small Rossby number, the presence of stratification allows for surface-intensified flows that are decoupled from the topography. We will therefore focus on a class of flows that to the lowest order are aligned with the bottom topography throughout the water column, that is, the magnitude of the velocity can vary vertically, but the velocity vector is parallel to the isobaths at all depths; this is sometimes referred to as equivalently barotropic flow (e.g., Killworth 1992). We will first make a Rossby number expansion of the stratified equations to describe the flow variations along curving isobaths. We then briefly illustrate that similar results can be obtained also in the quasigeostrophic framework in a limit of weak stratification and moderate bottom slopes. To the lowest order in the Rossby number w0 is zero, which implies that the buoyancy [Eq. (17)] to this order becomes u0 =b0 5 0 .

(43)

To satisfy Eq. (43), the horizontal gradient of b0 has to be parallel to that of f0. Here, we consider zeroth-order velocity fields that at all depths are aligned with the local isobaths, that is, f0 5 f0(H, z) and b0 5 b0(H, z). This choice automatically satisfies the bottom boundary condition [Eq. (31)]. Accordingly, there can be a vertical velocity shear tied to the horizontal variation of b0, but the shear is by our specific choice constrained to be locally parallel to the isobaths. As a consequence, the zeroth-order flow conserves the total volume transport between the isobaths. From Eqs. (20) and (21), it follows that potential vorticity conservation to the zeroth order is given by u0 =Q0 5 0, Q0 [ f

›b0 , ›z

(44)

stating that the vertical isopycnal spacing is constant along the isobaths. To the first order in the Rossby number, the buoyancy equation [Eq. (17)] becomes u0 =b1 1 u1 =b0 1 w1

›b0 5 0. ›z

ð0 z

u0 =z0 dz.

allow us to calculate the first-order pressure f1. Below, we present a method for solving this set of equations. An interpretation of the results in terms of potential vorticity conservation is given in section 4c. Following the ideas used to derive Eq. (39) in the homogeneous case, we solve the first-order buoyancy equation along the zeroth-order, isobath-following streamlines. The result is obtained by first noting that ð0 z

u0 =z0 dz 5 t =

ð0

›f0 z0 =f dz, z ›H

u1 =b0 5 2t =B1

=

›b0 f, ›H

(47)

where we again have used the scalar triple product rule. Using this, the buoyancy equation [Eq. (45)] can be rewritten as ›f1 ›f0 ›b 2 f1 0 2 F 5 0, t= ›z ›H ›H

(48)

where we have used the hydrostatic relation and introduced F(x, y, z) [

ju0 j2 ›b0 ›b0 1 2 ›H ›z

ð0

›f0 z0 dz. z ›H f

(49)

Using similar manipulations, the boundary condition for the vertical velocity at the bottom (i.e., z 5 2H) can be written as ð 1 0 ›f0 t = B1 1 z0 dz 5 0. f 2H ›H

(50)

Note that the quantities in the square brackets in Eqs. (48) and (50) are invariant along the depth contours. Equation (48) is a linear, first-order differential equation in z for f1, where the inhomogeneous term F represents the buoyancy advection due to the zeroth-order fields. An explicit solution will be presented after we have briefly commented on how quasigeostrophic versions of the results presented above can be obtained.

(45)

QUASIGEOSTROPHIC LIMIT

The bottom boundary condition involving w1 is given by Eq. (32). The vertical velocity w1 can be calculated from the vorticity equation [Eq. (27)] as w1 5 2f 21

2447

(46)

Since the vertical velocity w1 is determined by the zerothorder fields, Eq. (45) together with the hydrostatic relation [Eq. (3)] and the first-order momentum equation [Eq. (25)]

In the quasigeostrophic (QG) regime, the vertical buoyancy gradient is taken to be dominated by a background stratification, associated with the buoyancy frequency N(z), and b now represents a small perturbation on the background stratification (e.g., Allen 1991; Vallis 2006). A key parameter in this regime is the Burger number Bu [

N2H2 , f 2 L2

(51)

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VOLUME 46

where L is the horizontal length scale of the flow. The Burger number measures the ratio between vertical and horizontal buoyancy advection. As a result, the QG buoyancy equation becomes to the zeroth order in the Burger number identical to the zeroth-order Rossby number buoyancy equation [Eq. (43)] (see, e.g., Vallis 2006). In fact, also to the first order in a Burger number expansion of the QG equations, we can identify a very close correspondence with the first-order Rossby number equations presented above. To obtain this correspondence, we require that Ro Bu 1 and further that the bottom slope [Eq. (28)] is on the order of d ; Ro/Bu.

(52)

The usual QG assumption is that d ; Ro, which gives the boundary condition in Eq. (30). The value of d given by Eq. (52) is larger than Ro, though still taken to be small compared to one and yields in a Burger number expansion the zeroth- and first-order bottom boundary conditions in Eqs. (31) and (32). In the appendix, a simple, linear QG model is presented where d can be smoothly increased to yield the regime with strong topographic steering specified by Eq. (52). In short, the QG Burger number expansion counterparts of the relations presented above are similar, but we require that the isopycnal slopes of b0 are small and make the substitution ›b0 / N 2 (z) , ›z

(53)

neglect all the ju0j2 terms, for example, B1 / f1, and take u1 5 f21k 3 =f1. We also must replace the potential vorticity equation [Eq. (44)] with the QG version u0 =

b › f 02 5 0. ›z N

(54)

where s is an along-isobath coordinate and K0up denotes the upstream value. Accordingly, B1 5 f1 1 DK1. We also introduce the analogously defined along-isobath variation in relative vorticity Dz0. A particularly simple solution arises in the case where b0 is a function of z only, and db0/dz 5 N2 is a constant, that is, the zeroth-order flow is purely barotropic. In this special case, the first-order conservation of potential vorticity [cf. Eq. (20)] is simply given by ›b u0 = N 2 z0 1 f 1 5 0, ›z

c. A simple illustrative case To illustrate the dynamics encapsulated in Eqs. (48)– (50), we consider a situation where at some upstream location the topography and the flow is essentially oriented along a straight continental slope but where farther downstream the depth contours curve and/or diverge or converge. At the upstream location, we can without loss of generality set f1 5 0 and B1 5 0. For notational convenience, we introduce the along-isobath change in kinetic energy relative to the upstream location DK0 [ K0 (s, H, z) 2 K0up (H, z), K0 [ ju0 j2 /2,

FIG. 6. Theoretically calculated first-order surface height anomalies (cm) associated with the vorticity anomalies shown in Fig. 5, calculated from Eq. (60) using the following values: Hmax 5 3000 m, L 5 300 km, f 5 1024 s21 , NHmax 5 1.7 m s21, and 2Hmax/g(df0 /dH) 5 0.5 m. The white line illustrates the upstream velocity profile (see caption of Fig. 5). The white contours (21.5, 21.0, and 20.5 cm) show the contribution to the surface height anomaly due to the kinetic energy variations, which is included in the full anomaly indicated by the color shading. According to the theory, the first-order surface height anomalies should be equal to the along-isobath variations in the altimetrically derived ADT.

(55)

(56)

which can be recognized as steady QG potential vorticity conservation. The first-order solutions obtained from Eqs. (48)–(50) are f1 (s, H, z) 5

Dz0 df N 2 (H 2 2 z2 ) 2 DK0 , 2H 0 1 2 f dH (57)

Dz b1 (s, H, z) 5 2 0 N 2 z , f Dz0 df N2H3 C1 (s, H) 5 2 2H 2 0 1 , and f dH 3

(58) (59)

AUGUST 2016

h1 (s, H) 5

BROOMÉ AND NILSSON

Dz0 df N2H2 2H 0 1 2 DK0 /g. gf dH 2

(60)

Since the zeroth-order flow is barotropic in this case, these first-order fields are similar to the ones discussed in relation to Fig. 5; the quasigeostrophic versions result when DK0 is set to zero. The baroclinic component of the solution is related to the terms proportional to N2; if N2 is set to zero the purely barotropic case is recovered. Thus, as discussed in section 3, the free-surface height can be partitioned into a barotropic bottom pressure component and a steric height component. Here, the first-order variations in buoyancy and steric height result from vertical advection of the background stratification; downward vertical velocities, occurring where Dz0 is positive, result in positive buoyancy anomalies. Equation (56) shows that this is a consequence of potential vorticity conservation. Figure 6 illustrates Eq. (60), for which df0/dH is a constant and the slope becomes narrower halfway along the depth contours in the domain. Note that Eq. (60) is defined along depth contours, so that the anomalies in the solution are compared to upstream values at the same depth at x 5 0. The flow is essentially a jet concentrated to the steepest part of the slope, with cyclonic (anticyclonic) vorticity on the outer (inner) part of the slope; this yields the vorticity anomaly pattern shown in Fig. 5a, having a somewhat stronger anticyclonic anomaly. The surface height anomaly in Fig. 6, on the other hand, has a larger positive than negative amplitude, which is a result of the H-dependent terms in Eq. (60). To estimate possible magnitudes of the along-isobath surface height changes, we have chosen parameters broadly representative of the Norwegian Atlantic Slope Current (see the caption of Fig. 6). The slope j=Hj ranges from about 2% to 4% along the coast, a variation that is even larger along the NwASC. In the resulting surface height anomaly, the changes in kinetic energy play a minor role except for near the coast. Thus, where a continental slope narrows, one typically expects to see the surface height increase along the isobaths on the outer slope. On the inner slope, the surface height can be expected to decrease but with a magnitude that is smaller than the surface height changes on the outer slope. Along the Norwegian coast, the buoyancy decreases offshore, that is, ›b0/›H , 0. As discussed in the appendix, this amplifies the surface height anomaly h1. From Fig. 5, it is clear that where Dz0 is positive, the corrected streamlines are displaced toward deeper water. Since ›b0/›H , 0, this is associated with an offshore transport of light water, creating an additional positive

2449

buoyancy anomaly that is in phase with the one arising from the vertical advection.

5. Intercomparison of model and observations According to the model presented in the sections above, along-slope variations in relative vorticity and kinetic energy give rise to variations in sea surface height. To examine this in the observational data, the relative vorticity is computed from the ADT according to z 5 g=2hADT/f. Figure 7 shows the ratio of this and the Coriolis parameter f, representing the Rossby number. It can be seen to trace out the slope current as well as the frontal outer branch, with a negative (positive) value on the right (left) flank. It can also be noted that the relative vorticity in Fig. 7 displays an anticyclone in the LB (see previous mentioning of the Lofoten Basin eddy). A closer examination reveals an increase in amplitude around the two locations of maxima in steric height (cf. Figs. 3 and 4), similar to the pattern of the positive vorticity anomaly on the deep flank of the current and negative vorticity anomaly on the shallow flank in Fig. 5a. From Fig. 7 the along-isobath vorticity variations are estimated to be Dz/f ’ 0.1. To make an estimate of the associated surface height anomaly expected from Eq. (60), we assume numbers typical for the NwASC: df0 Dh 0:1 ; 10 ; 0:002 m s22 ;g DH 500 dH H ; 700 m

(61)

NH ; 1:7 m s21 . The resulting surface height anomaly is 3 cm, which is the same order of magnitude as the along-slope variations in the ADT (Fig. 4). To further compare theory and observations, we simulate the along-slope first-order correction to the free-surface height, h1 in Eq. (60), by interpolating the ADT to the 700-m depth contour, which lies on the continental slope all along the NwASC. The result is shown in Fig. 8, together with the gradient (slope) and the Laplacian (curvature) of the topography, interpolated to the same isobath. We find a correlation of the ADT with both the slope and the Laplacian, with correlation coefficients of 0.6 and 0.7, respectively; thus, the observed surface height anomalies along the slope are located where the topography steepens and/or curves, as predicted by the model. The maxima in the gradient and the Laplacian are close to, but not exactly, collocated and thus reinforce each other’s effect on the vorticity of the flow to some extent. A geostrophic approach of the ADT

2450


df Dhb /Dhs 5 H o dH

FIG. 7. The relative vorticity z divided by the Coriolis parameter f to estimate the Rossby number. The relative vorticity is given by the ADT [(z/f ) 5 (g/f )=2hADT]. Outlined are the contours of bottom depths of 500, 750, 1000, 1500, 2000, and 3000 m.

in Fig. 8 suggests a cross-contour flow: out from the coast where the ADT increases along the slope and toward the coast where it decreases. In other words, to compensate for the variations in relative vorticity due to the topographic features, the geostrophic current moves, in this case, toward deeper water (cf. Fig. 5b). For a more quantitative measure, the magnitude of the along-slope variations in the Laplacian of the topography is estimated to be 3 3 1027 m21. Together with the numbers in Eq. (61), Eq. (60) with Eq. (42) implies a sea surface height variation of about 2 cm in response to the observed along-slope variations in the curvature. The steepness of the topography similarly adds a few centimeters to this estimate. This gives a surface elevation of the same order of magnitude as the observed variations in ADT. We conclude that the model seems to predict several aspects of what we observe, specifically the surface height increases on the outer slope along the isobaths where the slope steepens and curves. However, there are also some discrepancies. For instance, a quick estimate of the predicted relative sizes of the barotropic and baroclinic contributions to the surface height variations, given by estimating the following quotient [cf. Eq. (60)]

VOLUME 46

2 2 N H 2

(62)

with typical numbers for the slope, see Eq. (61), predicts comparable sizes of the two terms. The observations, on the other hand, show smaller amplitude of the variations in bottom pressure than in steric height (see Fig. 4). This could be the result of a process not included in the model, for example, buoyancy gradients or eddies, that is amplifying the response of the steric height. The choice of depth contour for the along-slope representation of the data also affects the structure and correlation to some degree. For example, the correlation of the ADT with the Laplacian does not seem very sensitive to the choice of depth contour, keeping about 0.7 for the 500- and 900-m contours as well, while the correlation with the steepness of the slope is more sensitive and exhibits a slightly higher correlation on the 900-m contour and a weak (,0.4) correlation coefficient on the 500-m contour. Another discrepancy between the model and observations is the predicted negative anomaly in surface height on the inshore side of the current in Fig. 6, which is not found in observations. Some simplifications are made in Figs. 5 and 6, for example, df0/dH is considered as a constant and only the curvature term is taken into account, so we expect some discrepancies. We can also note that the response of the surface height to the vorticity anomaly in Fig. 5a is amplified on the deeper contours because of the H-dependent terms in Eq. (60). This is further amplified if horizontal buoyancy gradients are accounted for (see appendix). We can thus expect a positive surface anomaly on the deep flank to be the dominant response.

6. Summary and discussion In this study, we examine the dynamics of the NwASC, a cyclonic boundary current in the Nordic seas, with satellite altimetry as well as hydrographic data. First, we make a decomposition of the time-mean dynamic sea surface height as observed by satellites into a steric height and a dynamic bottom pressure, where the latter ideally should capture the topographic steering of the NwASC along the Norwegian continental slope. The presented bottom component essentially exhibits the anticipated topographic alignment along the slope, but regional exceptions are also present. Second, we further examine the topographic steering of the NwASC. Analysis of the time-mean altimetric and hydrographic data indicates dynamically nonnegligible variations and even local maxima in surface pressure

AUGUST 2016

BROOMÉ AND NILSSON

2451

FIG. 8. Along-slope variations in ADT (black, m) and the gradient (or slope, red) and Laplacian (or curvature, blue, m21) of the topography. The data are shown interpolated to the 700-m depth contour. Thick lines are the 50-pt running means. Correlations are between the running means of the ADT and the gradient (r 5 0.62) and Laplacian (r 5 0.71). The Laplacian and the gradient have been smoothed with Gaussian filters (with widths larger than the baroclinic Rossby radius) before being interpolated to the depth contour. The signs of the gradient and Laplacian are here set to match the sign of the corresponding vorticity anomaly according to the model equations.

and depth-integrated buoyancy along the isobaths in the core of the NwASC. This is somewhat remarkable since stationary topographic waves are not supported in cyclonic flows. To seek an explanation of these variations, we develop a simple adiabatic steady-state model of a cyclonic stratified boundary current over a topographic slope in the limit of small Rossby numbers. To the lowest order, the flow follows the isobaths, implying that changes in the along-stream topography cause changes in velocity and relative vorticity. Conservation of potential vorticity requires cross-isobath excursions, associated with vertical stretching or compression of fluid elements, to compensate for the changes in relative vorticity. The result is that along-stream topographic features that cause a local increase (decrease) in relative vorticity are associated with positive (negative) anomalies in bottom pressure and steric height and thus also in sea surface height. The model predicts along-isobath variations of a few centimeters along the NwASC and seems to qualitatively fit the observational data. The observational analysis has been limited to the Nordic seas and the NwASC, but the theoretical ideas are directly transferable to cyclonic boundary currents in other subpolar basins such as the Labrador Sea.

It should be noted, however, that the along-stream variations in depth-integrated buoyancy and sea surface height could be related to horizontal eddy advection and diabatic processes such as air–sea buoyancy fluxes rather than to vertical mean advection related to adiabatic potential vorticity conservation, as suggested here. To discuss this we write the depth-integrated buoyancy advection as ð0 2H

u =b dz 5

ð0 2H

[2N 2 w 2 = (u0 b0 )] dz 2 Q ,

(63)

where Q is the surface buoyancy flux, the overbar denotes a time mean, and the prime denotes time-dependent eddies. Assuming that the horizontal velocity u is largely along depth contours, the left-hand side of the equation essentially represents the along-isobath advection of steric height. The time-mean advection can thus be balanced by the mean vertical advection, the eddy flux divergence, and the air–sea buoyancy flux. Now, the air–sea buoyancy flux has the same sign over the whole NwASC (Segtnan et al. 2011; Isachsen et al. 2007), acting to densify water all along the current as the warm water loses heat to the atmosphere. Thus, the surface flux cannot account for an along-stream increase in buoyancy or steric height.

2452


Baroclinic eddies play an important role in the heat balance of the NwASC by fluxing heat from its core into a pool of water with lower mean advection between the two main branches of the Norwegian Atlantic Current (e.g., Isachsen et al. 2012). The intensity of the eddy field varies along stream in the NwASC, possibly reflecting varying degrees of baroclinic instability as the mean current speed and the topographic slope change (Spall 2010; Isachsen 2011; Poulin and Flierl 2005). Isachsen et al. (2012) estimate heat flux divergences in the NwASC due to the horizontal mean flow and eddies using surface drifter data, satellite data, and an eddy-permitting ocean circulation model. The heat flux divergence is examined along a path of the NwASC defined by its surface temperature maximum, which differs slightly between the data and the model simulation and does not strictly follow an isobath. A main result emerging both from the observational data and the model simulation is that the horizontal mean flow advection warms the current core at a rate that is comparable to the cooling due to the eddy flux divergence: the air–sea heat flux plays a secondary role in the heat budget. Thus, it is plausible that alongisobath variations in the eddy fluxes can cause local steric height maxima. Indeed, the maxima in the present data analysis are encountered in regions with more intense eddy activity (see, e.g., Isachsen et al. 2012). However, the estimates of the eddy heat flux divergence along the NwASC presented by Isachsen et al. (2012) are monotonic, acting to reduce the buoyancy. Thus, their study suggests that the eddy fluxes are not the primary mechanism for creating the along-isobath steric height maxima in the NwASC. As mentioned in section 3a, the gridded altimetric and hydrographic data have limited spatial resolution and could give partly erroneous impressions of alongstream variations. However, comparable variations are found in satellite-based ADT as well as in the two hydrographic datasets (WOA13 and the isopycnal climatology), which gives some confidence in the findings. In summary, the presented theoretical model in combination with the qualitative observational analysis of the time-mean circulation in the eastern Nordic seas indicate that the stationary anomalies in sea surface height and steric height along the NwASC can result from cross-isobath flow, governed by conservation of potential vorticity. As a concluding note, the Lofoten Basin is a subarctic hot spot in many ways, and numerous studies have focused on its rather spectacular heat content (cf. the steric height in Fig. 3) and the semipermanent

VOLUME 46

anticyclone residing in the deep basin and its possible connection to the NwASC (e.g., Volkov et al. 2013; Rossby et al. 2009; Orvik 2004). Several studies have suggested that anticyclonic eddies, which shed off from the slope current off the Lofoten Islands and drift westward, supply the heat to the central basin and maintain the semipermanent anticyclonic vortex (e.g., Köhl 2007; Raj et al. 2015). However, it is interesting to note that the present theory predicts that the hydrographic structure along a slope current should become deeper where the topographic slope increases. The simple example illustrated in Fig. 6 has some broad similarities with the hydrographic signature in the Lofoten Basin. It is an open question whether the hydrographic deepening occurring in the present simple model can have some seeding role in the complex chain of events that appear to control the dynamics of the Lofoten Basin. Acknowledgments. This work was supported by a grant from the Swedish National Space Board. The authors also thank Jonas Nycander for interesting discussions as well as two anonymous reviewers for valuable comments and suggestions.

APPENDIX A Quasigeostrophic Linear Stratified Model Here, we analyze the linear dynamics of a stratified steady flow over a gently sloping bottom using the quasigeostrophic equations. The purpose is to illustrate a few dynamical issues mentioned briefly in the main text, including some differences between steady flows that have shallow water to the right and left, dynamical lowpass filtering of the bottom topography, and how the flow becomes topographically steered when the Burger number Bu [Eq. (51)] is small and the bottom slope moderate. At the end, we discuss how vertical shear acts to magnify the surface pressure response to downstream changes in steepness or curvature of the isobaths. The problem is essentially the one studied by Rhines (1970), but we focus on the steady-state flow and allow for a mean flow with constant vertical shear. The basicstate mean flow, say u, is u(z) 5 U0 1 L(z 1 H0 );

›u 1 ›B 52 [ L, ›z f ›y

(A1)

where U0 and L are constants, and B 5 B(y) is the basicstate horizontal buoyancy. The vorticity and the buoyancy equations, governing the deviation from the background state, are given by

BROOMÉ AND NILSSON

AUGUST 2016

u u

›z ›w 5f , and ›x ›z

(A2)

›b ›B 1y 1 wN 2 5 0. ›x ›y

(A3)

The bottom depth is defined by H(x, y) 5 H0 1 Sy 2 h(x, y),

(A4)

where h is a small anomaly on a linearly sloping bottom. A key assumption is that the deviations from the mean depth H0 are small enough that the bottom boundary condition can be evaluated at z 5 2H0. The boundary conditions on the vertical velocity at the surface and the bottom are w(z 5 0) 5 0,

w(z 5 2H0 ) 5 2yS 1 u

›h , ›x

› f 2 ›2 f u =2 f 1 2 2 5 0 . ›x N ›z

(A6)

›f L 1 SN 2 /f , 2 l1 f 1 N 2 h 5 0, l1 [ ›z U0

(A7)

and the surface boundary condition can be written as ›f L 2 l2 f 5 0, l2 [ . ›z U0 1 LH0

(A8)

a. Vertical structure and the difference between cyclonic and anticyclonic mean flow By assuming a sinusoidal topographic perturbation h(x, y) } exp(ik x), where k is the horizontal wave vector [or formally taking the Fourier transform of h(x, y) and the equations], the solution to Eq. (A6) with its boundary conditions can be written as

A5

U0 (1 1 «)f h(k) f(k, 0) 5 , S 1 1 G2 k2

(A11)

which, using an inverse Fourier transform, corresponds to the potential vorticity–like equation U (1 1 «)f h(x, y) (G2 =2 2 1)f(x, y, 0) 5 2 0 S

(A12)

in x–y coordinates. Here, we have defined

By using Eq. (A3), the bottom boundary condition in Eq. (A5) can be written as

f(k, z) 5 A[cosh(xz) 1 (l2 /x) sinh(xz)],

where h(k) is the wavenumber representation of the topographic perturbation and x [ jkjN/f. The parameter (xH0)2 is a Burger number based on the horizontal scale L ; jkj21 of the topography h. When this parameter is large, the response is bottom trapped with a negligible surface pressure signature; the ratio between surface and bottom pressure is proportional to exp(2xH0). We are here interested in the opposite limit where xH0 1, that is, when the horizontal scale of the topographic perturbation is larger than the baroclinic Rossby radius. By expanding the denominator in Eq. (A10), retaining terms up to the second order in xH0, we obtain, after some simplifications, an expression for the surface pressure in this limit

(A5)

where we make the rigid-lid approximation and neglect the small term u =h. By using the hydrostatic and geostrophic relations, the equations and boundary conditions can be formulated in terms of the pressure f alone. Combining Eqs. (A2) and (A3) and taking N as constant yields the potential vorticity equation

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and

(A9)

N 2 h(k) , (x 2 l1 l2 /x) sinh(xH0 ) 1 (l1 2 l2 ) cosh(xH0 ) (A10)

H0 U0 «2 N 2 H02 1 « G [ 1 , 11«1 1 2 2 3 Sf 3 f 2

« [ LH0 /U0 ,

(A13)

where « is the ratio between the basic-state vertical velocity shear and the bottom velocity. In the definition of G2 , we recognize two length scales: the wavelength of stationary barotropic topographic waves in an anticyclonic mean flow [(H 0 jU 0 j)/(Sf )] 1/2 and the baroclinic Rossby radius NH 0 /f. The sign of G2 determines the character of the solutions to Eqs. (A11) and (A12). When G2 , 0, the solutions are wavelike and of hyperbolic character; in this case, Eq. (A12) describes forced steady topographic Rossby waves [see Holton (2004, section 7.7) for a discussion of the barotropic case]. On the other hand, when G2 . 0, the solutions are trapped to the topographic perturbations and have elliptic character. In the barotropic limit, that is, when N 5 0 and « 5 0, the sign of U 0 alone determines whether the solutions are of hyperbolic (U 0 , 0, anticyclonic mean flow having the shallow water to the left) or elliptic character (U 0 . 0, cyclonic mean flow having the shallow water to the right).

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b. Topographic steering and low-pass filtering at small Bu For cyclonic mean flows, the features of Eq. (A11) imply that the surface pressure response is a lowpassed version of the topographic perturbation (LaCasce et al. 2008), in which spatial scales that are smaller than about G are filtered out. In this limit, where (G/L) 2 1, the dominant balance in Eq. (A12) is z } 2h, that is, anomalies in relative vorticity balance water depth changes. Here, f(x, y, 0) becomes small, implying that the net flow approximately is the mean flow u, which passes over the topographic anomalies h rather than around them. In the opposite regime, where (G/L) 2 1, the dominant balance in Eq. (A12) is f(x, y, 0) } h, that is, the surface pressure and water depth anomalies are aligned and the flow follows the depth contours. If « ; 1 or smaller, we have

f1 (H, z 5 0) 5

Ro 1 Bu, d

Ro [

U0 (NH0 )2 SL , d [ , Bu [ . H0 fL ( fL)2 (A14)

As we have already assumed that Bu 1, (G/L)2 will be small if Ro/d 1; in the quasigeostrophic analysis in section 4, we required that Ro/d ; Bu. Thus, to the zeroth order in Bu, Eq. (A12) gives f(x, y, 0) 5 hfU0(1 1 «)/S. Combining this with the pressure associated with the mean flow u, which is 2yU0(1 1 «)/f, we get an isobath-aligned ‘‘zeroth-order’’ surface pressure field: f0 (H, z 5 0) 5

U0 (1 1 «)f (h 2 Sy) . S

›f0 (H, z) 5 2FH A(z) , ›H (A17)

where g and N2 are positive constants, and we have defined ›f (H, z 5 2H) , FH [ 2 0 ›H A(z) [ 1 1 s(z 1 H), s [ g/FH .

(A18)

Here, FH and s are positive constants, and sH is the ratio between the bottom-to-top baroclinic shear in velocity and the bottom velocity. Some calculations show that the first-order surface pressure field is ( " # Dz0 (x, y, 0) (sH)2 HFH 1 1 sH 1 f1 (x, y, 0) 5 f 3 ) N2H2 2sH , 11 1 3 2 (A19)

(A15)

To the first order in Bu, Eq. (A12) gives f1(x, y, 0) 5 G2=2hfU0(1 1 «)/S, or equivalently if we use Eq. (A15),

=2 f0 (H, z 5 0) H0 U0 f «2 N 2 H02 2« . 1 1 « 1 1 1 1 3 f2 S 3 2

We can compare these results with a solution to Eq. (48) that have analogous zeroth-order buoyancy and pressure fields: b0 (H, z) 5 2gH 1 N 2 z,

(G/L)2 ;

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(A16)

where we for simplicity have omitted terms involving the zeroth-order kinetic energy. By noting that (H0U0f)/S corresponds to HFH and that « corresponds to sH, we see that this first-order surface pressure field is essentially identical to the one given in Eq. (A16). The main difference in the underlying assumptions behind Eqs. (A16) and (A19) is that the latter is formally valid for order one variations in the depth H. Yet, it is important to note that they yield the same parametric dependence on the basic-state parameters, suggesting that quasigeostrophic theory gives qualitatively correct results, even when the bottom slope becomes steep. Finally, we note that a positive vertical shear in the zeroth-order flow, which is measured here by « or sH, amplifies the first-order surface pressure response [see Eqs. (A16) and (A19)]. This can be interpreted in terms of potential vorticity conservation, as the horizontal buoyancy gradients modify the potential vorticity gradients on the boundaries (Bretherton 1966). Alternatively, the amplification can be seen as a result of buoyancy advection, which brings lighter water offshore where the relative vorticity is anomalously positive. Using the near-surface horizontal buoyancy gradients along the Norwegian Atlantic Slope Current, we find s21 is on the order of 1000 m; that is, along the 1000-m isobath sH should be on the order of unity. Thus, the effects of the horizontal buoyancy gradient should be

BROOMÉ AND NILSSON

AUGUST 2016

of leading-order importance for the amplitude in the along-isobath pressure variations.

REFERENCES Aaboe, S., and O. A. Nøst, 2008: A diagnostic model of the Nordic Seas and Arctic Ocean circulation: Quantifying the effects of a variable bottom density along a sloping topography. J. Phys. Oceanogr., 38, 2685–2703, doi:10.1175/2008JPO3862.1. Allen, J. S., 1991: Balance equations based on momentum equations with global invariants of potential enstrophy and energy. J. Phys. Oceanogr., 21, 265–276, doi:10.1175/1520-0485(1991)021,0265: BEBOME.2.0.CO;2. Becker, J. J., and Coauthors, 2009: Global bathymetry and elevation data at 30 arc seconds resolution: SRTM30_PLUS, revised for marine geodesy, January 20, 2009. Mar. Geod., 32, 355–371, doi:10.1080/01490410903297766. Björk, G., and Coauthors, 2010: Flow of Canadian basin deep water in the western Eurasian Basin of the Arctic Ocean. Deep-Sea Res. I, 57, 577–586, doi:10.1016/j.dsr.2010.01.006. Böning, C. W., W. Bryan, W. R. Holland, and R. Döscher, 1996: Deep-water formation and meridional overturning in a highresolution model of the North Atlantic. J. Phys. Oceanogr., 26, 1142–1164, doi:10.1175/1520-0485(1996)026,1142: DWFAMO.2.0.CO;2. Bretherton, F. P., 1966: Critical layer instability in baroclinic flows. Quart. J. Roy. Meteor. Soc., 92, 325–334, doi:10.1002/ qj.49709239302. Carton, J. A., G. A. Chepurin, J. Reagan, and S. Häkkinen, 2011: Interannual to decadal variability of Atlantic Water in the Nordic and adjacent seas. J. Geophys. Res., 116, C11035, doi:10.1029/2011JC007102. Chafik, L., 2014: Dynamics and variability of the circulation in the North-Atlantic subpolar seas. Ph.D. thesis, Stockholm University, 38 pp. ——, J. Nilsson, Ø. Skagseth, and P. Lundberg, 2015: On the flow of Atlantic water and temperature anomalies in the Nordic Seas toward the Arctic Ocean. J. Geophys. Res. Oceans, 120, 7897– 7918, doi:10.1002/2015JC011012. Döös, K., and D. J. Webb, 1994: The Deacon cell and the other meridional cells of the Southern Ocean. J. Phys. Oceanogr., 24, 429–442, doi:10.1175/1520-0485(1994)024,0429: TDCATO.2.0.CO;2. Eldevik, T., J. E. Ø. Nilsen, D. Iovino, K. A. Olsson, A. B. Sandø, and H. Drange, 2009: Observed sources and variability of Nordic seas overflow. Nat. Geosci., 2, 406–410, doi:10.1038/ ngeo518. Fofonoff, N., 1962: Dynamics of ocean currents. Physical Oceanography, M. N. Hill, Ed., The Sea—Ideas and Observations on Progress in the Study of the Seas, Vol. 1, 323–395. Greenspan, H. P., 1968: The Theory of Rotating Fluids. 1st ed. Cambridge University Press, 327 pp. Helland-Hansen, B., and F. Nansen, 1909: The Norwegian Sea—Its Physical Oceanography Based upon the Norwegian Researches 1900-1904. Report on Norwegian Fishery and Marine Investigations Series, Vol. 2, Fiskeridirektoratets havforskningsinstitutt, 422 pp. Holton, J. R., 2004: An Introduction to Dynamic Meteorology. 4th ed. Elsevier Academic Press, 535 pp. Isachsen, P. E., 2011: Baroclinic instability and eddy tracer transport across sloping bottom topography: How well does a modified Eady model do in primitive equation

2455

simulations? Ocean Modell., 39, 183–199, doi:10.1016/ j.ocemod.2010.09.007. ——, and O. A. Nøst, 2012: The air-sea transformation and residual overturning circulation within the Nordic Seas. J. Mar. Res., 70, 31–68, doi:10.1357/002224012800502372. ——, C. Mauritzen, and H. Svendsen, 2007: Dense water formation in the Nordic Seas diagnosed from sea surface buoyancy fluxes. Deep-Sea Res. I, 54, 22–41, doi:10.1016/ j.dsr.2006.09.008. ——, I. Koszalka, and J. LaCasce, 2012: Observed and modelled surface eddy heat fluxes in the eastern Nordic Seas. J. Geophys. Res., 117, C08020, doi:10.1029/2012JC007935. Jakobsen, P. K., M. H. Ribergaard, D. Quadfasel, T. Schmith, and C. W. Hughes, 2003: Near-surface circulation in the northern North Atlantic as inferred from Lagrangian drifters: Variability from the mesoscale to interannual. J. Geophys. Res., 108, 3251, doi:10.1029/2002JC001554. Johannessen, J., and Coauthors, 2014: Toward improved estimation of the dynamic topography and ocean circulation in the high latitude and Arctic Ocean: The importance of GOCE. Surv. Geophys., 35, 661–679, doi:10.1007/s10712-013-9270-y. Killworth, P. D., 1992: An equivalent-barotropic mode in the Fine Resolution Antarctic Model. J. Phys. Oceanogr., 22, 1379–1387, doi:10.1175/1520-0485(1992)022,1379:AEBMIT.2.0.CO;2. Köhl, A., 2007: Generation and stability of a quasi-permanent vortex in the Lofoten Basin. J. Phys. Oceanogr., 37, 2637–2651, doi:10.1175/2007JPO3694.1. Koszalka, I., J. LaCasce, M. Andersson, K. Orvik, and C. Mauritzen, 2011: Surface circulation in the Nordic Seas from clustered drifters. Deep-Sea Res. I, 58, 468–485, doi:10.1016/ j.dsr.2011.01.007. LaCasce, J., O. Nøst, and P. Isachsen, 2008: Asymmetry of free circulations in closed ocean gyres. J. Phys. Oceanogr., 38, 517–526, doi:10.1175/2007JPO3789.1. Li, S., and T. A. McClimans, 2000: On the stability of barotropic prograde and retrograde jets along a bottom slope. J. Geophys. Res., 105, 8847–8855, doi:10.1029/1999JC000006. Locarnini, R. A., and Coauthors, 2013: Temperature. Vol. 1, World Ocean Atlas 2013, NOAA Atlas NESDIS 73, 40 pp. Marshall, D. P., and C. E. Tansley, 2001: An implicit formula for boundary current separation. J. Phys. Oceanogr., 31, 1633–1638, doi:10.1175/1520-0485(2001)031,1633:AIFFBC.2.0.CO;2. Marshall, J., and K. Speer, 2012: Closure of the meridional overturning circulation through Southern Ocean upwelling. Nat. Geosci., 5, 171–180, doi:10.1038/ngeo1391. Mauritzen, C., 1996: Production of dense overflow waters feeding the North Atlantic across the Greenland-Scotland Ridge. Part 1: Evidence for a revised circulation scheme. Deep-Sea Res. I, 43, 769–806, doi:10.1016/0967-0637(96)00037-4. Mork, K. A., and Ø. Skagseth, 2010: A quantitative description of the Norwegian Atlantic Current by combining altimetry and hydrography. Ocean Sci., 6, 901–911, doi:10.5194/ os-6-901-2010. Nøst, O. A., and P. E. Isachsen, 2003: The large-scale time-mean ocean circulation in the Nordic Seas and Arctic Ocean estimated from simplified dynamics. J. Mar. Res., 61, 175–210, doi:10.1357/002224003322005069. ——, J. Nilsson, and J. Nycander, 2008: On the asymmetry between cyclonic and anticyclonic flow in basins with sloping boundaries. J. Phys. Oceanogr., 38, 771–787, doi:10.1175/ 2007JPO3714.1. Orvik, K. A., 2004: The deepening of the Atlantic water in the Lofoten Basin of the Norwegian Sea, demonstrated by using

2456


an active reduced gravity model. Geophys. Res. Lett., 31, L01306, doi:10.1029/2003GL018687. ——, and P. Niiler, 2002: Major pathways of Atlantic water in the northern North Atlantic and Nordic Seas toward Arctic. Geophys. Res. Lett., 29, 1896, doi:10.1029/2002GL015002. Poulain, P.-M., A. Warn-Varnas, and P. Niiler, 1996: Near-surface circulation of the Nordic seas as measured by Lagrangian drifters. J. Geophys. Res., 101, 18 237–18 258, doi:10.1029/ 96JC00506. Poulin, F., and G. Flierl, 2005: The influence of topography on the stability of jets. J. Phys. Oceanogr., 35, 811–825, doi:10.1175/ JPO2719.1. Raj, R. P., 2013: The circulation of the Norwegian Sea—An investigation from space and ocean. Ph.D. thesis, University of Bergen, 175 pp. ——, L. Chafik, J. E. Ø. Nilsen, T. Eldevik, and I. Halo, 2015: The Lofoten vortex of the Nordic Seas. Deep-Sea Res. I, 96, 1–14, doi:10.1016/j.dsr.2014.10.011. Rhines, P., 1970: Edge-, bottom-, and Rossby waves in a rotating stratified fluid. Geophys. Astrophys. Fluid Dyn., 1, 273–302, doi:10.1080/03091927009365776. Rio, M., S. Guinehut, and G. Larnicol, 2011: New CNES-CLS09 global mean dynamic topography computed from the combination of GRACE data, altimetry, and in situ measurements. J. Geophys. Res., 116, C07018, doi:10.1029/ 2010JC006505. Rossby, T., V. Ozhigin, V. Ivshin, and S. Bacon, 2009: An isopycnal view of the Nordic Seas hydrography with focus on properties of the Lofoten Basin. Deep-Sea Res. I, 56, 1955–1971, doi:10.1016/j.dsr.2009.07.005. Rudels, B., R. D. Muench, J. Gunn, U. Schauer, and H. J. Friedrich, 2000: Evolution of the Arctic Ocean boundary current north of the Siberian shelves. J. Mar. Syst., 25, 77–99, doi:10.1016/ S0924-7963(00)00009-9. Schlichtholz, P., 2007: Density-dependent variations of the alongisobath flow in the East Greenland Current from the Fram Strait to the Denmark Strait. J. Geophys. Res., 112, C12022, doi:10.1029/2006JC003987. Segtnan, O., T. Furevik, and A. Jenkins, 2011: Heat and freshwater budgets of the Nordic Seas computed from atmospheric

VOLUME 46

reanalysis and ocean observations. J. Geophys. Res., 116, C11003, doi:10.1029/2011JC006939. Skagseth, Ø., K. A. Orvik, and T. Furevik, 2004: Coherent variability of the Norwegian Atlantic Slope Current derived from TOPEX/ERS altimeter data. Geophys. Res. Lett., 31, L14304, doi:10.1029/2004GL020057. Søiland, H., and T. Rossby, 2013: On the structure of the Lofoten Basin eddy. J. Geophys. Res. Oceans, 118, 4201–4212, doi:10.1002/jgrc.20301. ——, M. Prater, and T. Rossby, 2008: Rigid topographic control of currents in the Nordic Seas. Geophys. Res. Lett., 35, L18607, doi:10.1029/2008GL034846. Spall, M. A., 2005: Buoyancy-forced circulations in shallow marginal seas. J. Mar. Res., 63, 729–752, doi:10.1357/ 0022240054663204. ——, 2010: Non-local topographic influences on deep convection: An idealized model for the Nordic Seas. Ocean Modell., 32, 72–85, doi:10.1016/j.ocemod.2009.10.009. ——, and R. S. Pickart, 2001: Where does dense water sink? A subpolar gyre example. J. Phys. Oceanogr., 31, 810–826, doi:10.1175/ 1520-0485(2001)031,0810:WDDWSA.2.0.CO;2. Sverdrup, H. U., 1947: Wind-driven currents in a baroclinic ocean: With application to the equatorial currents of the eastern Pacific. Proc. Natl. Acad. Sci. USA, 33, 318–326, doi:10.1073/pnas.33.11.318. Vallis, G. K., 2006: Atmospheric and Oceanic Fluid Dynamics. Cambridge University Press, 745 pp. Volkov, D. L., T. V. Belonenko, and V. R. Foux, 2013: Puzzling over the dynamics of the Lofoten Basin—A sub-Arctic hot spot of ocean variability. Geophys. Res. Lett., 40, 738–743, doi:10.1002/grl.50126. Walin, G., G. Broström, J. Nilsson, and O. Dahl, 2004: Baroclinic boundary currents with downstream decreasing buoyancy: A study of an idealized Nordic Seas system. J. Mar. Res., 62, 517–543, doi:10.1357/0022240041850048. Zhang, J., D. A. Rothrock, and M. Steele, 1998: Warming of the Arctic Ocean by a strengthened Atlantic inflow: Model results. Geophys. Res. Lett., 25, 1745–1748, doi:10.1029/ 98GL01299. Zweng, M., and Coauthors, 2013: Salinity. Vol. 2, World Ocean Atlas 2013, NOAA Atlas NESDIS 74, 39 pp.