OSOcean ScienceOSOcean Sci.1812-0792Copernicus GmbHGöttingen, Germany10.5194/os-10-967-2014The Rossby radius in the Arctic OceanNurserA. J. G.BaconS.s.bacon@noc.ac.ukhttps://orcid.org/0000-0002-2471-9373National Oceanography Centre, Southampton, UKS. Bacon (s.bacon@noc.ac.uk)28November201410696797525September201322October201310October201429October2014This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://www.ocean-sci.net/10/967/2014/os-10-967-2014.htmlThe full text article is available as a PDF file from https://www.ocean-sci.net/10/967/2014/os-10-967-2014.pdf
The first (and second) baroclinic deformation (or Rossby) radii are
presented north of ∼60∘ N, focusing on deep basins
and shelf seas in the high Arctic Ocean, the Nordic seas, Baffin Bay, Hudson
Bay and the Canadian Arctic Archipelago, derived from climatological ocean
data. In the high Arctic Ocean, the first Rossby radius increases from
∼5 km in the Nansen Basin to ∼15 km in the
central Canadian Basin. In the shelf seas and elsewhere, values are low
(1–7 km), reflecting weak density stratification, shallow water, or both.
Seasonality strongly impacts the Rossby radius only in shallow seas, where
winter homogenization of the water column can reduce it to below 1 km.
Greater detail is seen in the output from an ice–ocean general circulation
model, of higher resolution than the climatology. To assess the impact of
secular variability, 10 years (2003–2012) of hydrographic stations along
150∘ W in the Beaufort Gyre are also analysed. The first-mode
Rossby radius increases over this period by ∼20 %. Finally,
we review the observed scales of Arctic Ocean eddies.
Introduction
The first baroclinic Rossby radius of deformation is of fundamental
importance in atmosphere–ocean dynamics. It is the horizontal scale at which
rotation effects become as important as buoyancy effects, or more precisely,
the horizontal scale of perturbation over which the vortex stretching and
relative vorticity associated with sloping isopycnals make approximately
equal contributions to potential vorticity. It is the horizontal scale of
geostrophic relaxation, and the “natural” scale of baroclinic boundary
currents, eddies and fronts. It sets the scale of the waves that grow most
rapidly as a result of baroclinic instability, and the long Rossby wave
speed (Gill, 1982; Chelton et al., 1998; Saenko, 2006).
In the context of ocean models, it is important to know the field of the
Rossby radius so that we know where models will describe boundary currents
and the eddy field adequately and where they will not. A minimum of two
grid points per eddy radius is necessary to resolve eddies adequately, and
one grid point per radius to “permit” them: e.g. Smith et al. (2000), Hecht
and Smith (2008). Hallberg (2013) suggests that eddy parameterizations may
no longer be necessary once the ratio of the baroclinic deformation radius
to a model's effective grid spacing is greater than a value of about 2,
where “effective spacing” means the grid-diagonal distance. The typical
best resolution in oceanic general circulation models (OGCMs) is currently
∼0.1∘ (ca. 10 km). In the deeper waters of the Arctic
Ocean and over much of the Nordic seas, stratification is generally weak, so
the Rossby radius might be expected to be significantly smaller than the
30–50 km characteristic of the mid-latitude oceans. OGCMs will therefore
(typically) be eddy-permitting in the Arctic region at best. Over the broad
Arctic Ocean shelf seas, the Rossby radius will be even smaller, and here
OGCMs will not even be eddy-permitting. Chelton et al. (1998) – C98
hereafter – describe the quasi-global geographical variability of the first
baroclinic Rossby radius, but their analysis does not extend north of
∼60∘ N. This motivates the present study, which is
structured as follows. In Sect. 2, we present our methods and data;
Sect. 3 describes results, including uncertainties; Sect. 4 is
discussion; and Sect. 5 presents some final remarks.
Methods and data
The standard method of finding the internal deformation (Rossby) radii
involves solving the linearized quasi-geostrophic potential vorticity
equation for zero background mean flow, and is described by C98. Briefly,
solutions for the velocity are separated into horizontal and vertical
components, where the structure of the vertical velocity φ(z) must
satisfy
N-2(z)d2φdz2+c-2φ=0,
where z is the vertical coordinate, and N is the buoyancy frequency, defined
by
N-2=-gρ∂ρloc∂z=gα∂θ∂z-β∂S∂z.
Here g is the acceleration due to gravity, ρ is in situ density, and
ρloc is the potential density referenced to the depth at which
N2 is being calculated. In the term on the RHS (an approximate form),
θ and S are potential temperature and salinity respectively, and
α and β are the thermal expansion and haline contraction
coefficients (Gill, 1982). The boundary conditions to be satisfied are zero
vertical velocity at the surface and ocean floor:
φ=0atz=-H,z=0,
where H is ocean depth. Solutions are possible only for certain values of
c-2 (the eigenvalues). The corresponding ci (decreasing with
increasing i) are then the phase speeds of the internal gravity waves, for
internal modes i=1,2,…, while the Rossby radii Ri are
Ri=ciciff,
where f=2Ω sinΘ is the Coriolis parameter for earth
rotation rate Ω and latitude Θ. The ci may be found
exactly by numerically integrating Eq. (1) to find values of c that satisfy
the conditions of Eq. (3). The eigenvalue problem of Eq. (1) depends only on
N2, so that solutions only require knowledge of the local vertical
density gradient. Variations in f have little influence on Ri within the
Arctic Ocean and Nordic seas – sin (90∘, 80∘, 70∘, 60∘) = (1, 0.985, 0.940, 0.866) – but are, of
course, important in reducing the deformation radius from low to high
latitudes.
In the presence of sloping topography, the separation of disturbances into
vertical and horizontal modes breaks down, because the vertical velocity
need not be zero on the ocean floor, so the conditions of Eq. (3) no longer
hold. Hence the idea of an internal-wave speed (and therefore Rossby radius)
that is independent of the horizontal structure is no longer strictly
appropriate. However, simple “flat-bottomed” values of the Rossby radius
(which we present below) remain useful estimates of the various horizontal
scales outlined in Sect. 1.
We also calculate the Rossby radius using the WKBJ/LG method, as derived by
C98, wherein approximate, analytical solutions to Eq. (1) are sought,
subject to the boundary conditions of Eq. (3). The ci are found by this
method to be
ci=πi-1∫-H0Ndz
subject to the requirement that φ should change slowly in the
vertical compared with the length scale (c/N). C98 found that this
approximation worked surprisingly well in generating a global climatology of
R1. Equation (5) can be further simplified, assuming constant N (Gill,
1982), which is useful for order-of-magnitude estimation of Rossby radii:
ci=NHiπ.
Bathymetry (m) of the Arctic Ocean and environs. The four major
basins of the deep Arctic Ocean are labelled: Canadian (C), Makarov (M),
Amundsen (A), Nansen (N). Representative station positions from the BGEP along 150∘ W are shown as white
dots.
We calculate the fields of the internal Rossby radius from the temperature
and salinity fields of the Polar Science Center Hydrographic Climatology
(PHC; Steele et al., 2001). PHC fields are provided on a 1∘× 1∘
latitude–longitude grid, with vertical resolution
increasing from 10 m near the surface, through 100 m at mid-depths (300–1500 m),
to 500 m below 2000 m. We use here the annual and seasonal fields, where
the seasons are defined in the PHC as the months of July, August and
September (summer), and March, April and May (winter).
To assess the magnitude of interannual variability of the Rossby radius, we
use hydrographic profile data from the Beaufort Gyre Exploration Project
(BGEP; Proshutinsky et al., 2009). We focus on the annually repeated section
(2003–2012) along 150∘ W, between the north Alaskan shelf
and the central Beaufort Gyre in the Canadian Basin. Station locations are
included in Fig. 1.
We also calculate the fields of the internal Rossby radius from the
temperature and salinity fields produced by the OCCAM global 1/12∘
model (Marsh et al., 2009). In the Arctic, the horizontal resolution is
∼9 km. The model has 66 levels in the vertical and includes
27 levels in the upper 400 m with thickness ranging from 5.4 m in the
uppermost layer to 48 m at 400 m and to 103 m at 1000 m. In the Arctic,
OCCAM was initialized with PHC. It was then run from 1985 to 2004 (Marsh et
al., 2009) using surface fluxes generated from bulk formulae using model sea
surface temperature and atmospheric output from the US National Centers for
Climate Research, together with satellite solar forcing and precipitation.
The OCCAM model bathymetry, which we use to illustrate Arctic Ocean regional
bathymetry (Fig. 1), is derived from the bathymetry of Smith and Sandwell (1997),
patched north of 72.0∘ N with the International
Bathymetric Chart of the Arctic Ocean (IBCAO) data set (Jakobsson et al.,
2000); see Aksenov et al. (2010a) for further details.
The usefulness of OCCAM in the Arctic has been demonstrated in a series of
recent papers: Aksenov et al. (2010a, b, 2011) describe the Atlantic
water inflows, the polar water outflows, and the representation of the
Arctic Circumpolar Boundary Current in the model. We use OCCAM output for a
number of reasons. The model has high horizontal and vertical resolution,
and realistic coastlines, so it interpolates and (to some extent)
extrapolates the climatological initialization. The model imposes dynamical
consistency throughout the domain, and by choosing output a few years from
the start of the run (here we inspect March and August 1992), most
inconsistencies associated with spin-up are avoided. The model is spun up
from rest, and Fig. 1 in Aksenov et al. (2010a) show that global mean
kinetic energy stabilizes after a few years, so that the model's dynamical
state is close to equilibrium, while the thermodynamic state has drifted
little from the initial conditions.
Results
The annual mean first and second mode Rossby radii (R1, R2) and the
mode 1 seasonal (summer minus winter) difference as derived from PHC fields
are shown in Fig. 2. Mean values for sub-regions are presented in Table 1,
and the sub-regions are defined in Table 1 and Fig. 3. There is
substantial variation in deformation radii over the Arctic. Considering
R1 first, in the deep basins of the Arctic Ocean, its annual mean values
increase quasi-monotonically, with typical values ranging from
∼7 km in the Nansen Basin, through 9–10 km in the Amundsen
Basin, 11 km in the Makarov Basin, to the largest values of ∼15 km, found towards the centre of the Canadian Basin. This reflects the
increase in stratification resulting from the progressive decrease in
upper-ocean salinity from the Amundsen to the Canadian Basin (e.g. Carmack,
2000).
(a) PHC annual mean Rossby radius (km), mode 1. (b) PHC Rossby radius (km), mode 1, seasonal difference (summer minus
winter). (c) PHC annual mean Rossby radius (km), mode 2.
Regional averages of Rossby radii mode 1 (summer, winter and
annual) and mode 2 (annual). The tabulated regions are shown in Fig. 3, to
which the key numbers refer. The Amerasian Basin combines the Canadian and
Makarov basins; the Eurasian Basin combines the Amundsen and Nansen basins.
RegionKeyMode 1 (km) Mode 2 (km)SummerWinterAnnualAnnualDeep Arctic Ocean Amerasian Basin111.211.011.15.2Eurasian Basin27.97.77.84.6Environs of Canadian waters Canadian Archipelago36.35.96.12.9Hudson Bay & Foxe Basin44.82.73.71.9Baffin Bay56.35.45.82.9Nordic seas Greenland Sea65.14.44.72.6Iceland Sea75.03.84.32.4Norwegian Sea87.26.56.93.0Eurasian Shelf seas Barents Sea93.31.92.51.3White Sea103.52.42.71.4Kara Sea114.63.03.81.9East Siberian Shelf seas123.72.53.11.4
The deep Nordic seas are divided by the mid-basin ridge system. The
Norwegian Sea contains the Atlantic-dominated inflows, where R1∼7 km. The Iceland and Greenland seas contain the
polar-dominated outflows, where R1∼3 km. All the shallow
shelf seas show very small values of R1 – generally less than 2 km, and
in places significantly less than 1 km. The seasonal variation is most
pronounced in the shallow shelf seas, where riverine freshwater inputs and
sea ice melt cause high summer stratification, and wintertime convective
homogenization of the water column causes low stratification.
Amplitudes of R2 are roughly half of those of R1, because the wave
speed solutions scale with mode number (see Eq. 5 below) and the contrast
between shallow shelf seas and deep ocean is similar. However, the mode 2
structure is subtly different to that of mode 1, in that the trans-polar
increasing tendency is largely absent.
Division of Arctic Ocean and adjacent regions for average Rossby
radius calculations; see Table 1 for identification of sub-regions by key
number.
There are of course substantial uncertainties associated with these
estimates. The first contribution to the uncertainty is secular (interannual
to decadal) variability, and the most notable secular variability in Arctic
Ocean properties is the increase in stored freshwater in the Beaufort Gyre
within the Canadian Basin (Rabe et al., 2011; Morison et al., 2012),
attributed to gyre spin-up (Giles et al., 2012). We calculate R1 along
the BGEP 150∘ W section for each of the 10 years
2003–2012, mainly between 72 and 79∘ N (see dots in Fig. 1). The section passes
through the centre of the gyre and also through the surface salinity
minimum, and therefore the stratification maximum (Fig. 4). As freshwater
storage increases, so too does R1 across the whole section, by
∼3 km overall, or ∼2 % per year.
Interestingly, there is an upward drift in R1 between 2003 and 2007, then a
jump of > 1 km from 2007 to 2008, after which values are
relatively stable. The jump coincides with the period of unusually low sea
ice extent of September 2007 (Stroeve et al., 2008). PHC results lie
generally within the lower, earlier range of BGEP values, but with evidence
of smoothing, on which we comment below.
Rossby radius mode 1 (km) along BGEP
150∘ W hydrographic section for 2003–2012; years are
identified by colour in the key. PHC values along 150∘ W are
shown in black, with grey shading showing the seasonal range.
Considering uncertainties in the vertical, there is no contribution from the
calculation of N2, which is exact, through use of locally referenced
potential density for both observations and model. However, a possible
contribution arises from the limited vertical resolution of the PHC data.
This is inspected by comparison of the Rossby radii derived from
full-resolution (1 dbar) BGEP profiles with results obtained by decimation
down to PHC resolution. Figure 5 shows the resulting differences (decimated
minus full), for which the mean difference is -0.05±0.07 km (1 SD),
or ∼0.5 %. The bias is a consequence of the tendency of
decimation to reduce the density gradient.
A “horizontal” contribution to uncertainty results from the smoothing due
to the gridding process used in generating the PHC climatology. At any depth
that is crossed by a density front which is included within the search
radius, horizontal averaging will increase the density on the “deep” side
and decrease the density on the “shallow” side (referring to isopycnal
depths), with consequent impacts on ∂ρ/∂z. C98 were able to form a regional
(North Atlantic) assessment of the significance of horizontal smoothing by
comparison of their results derived from their global horizontally averaged
database with those from a parallel, isopycnally averaged product. C98 found
a quasi-random uncertainty of 5 % (1 SD), plus a systematic and larger
contribution from averaging across the Gulf Stream and its extension, which
they illustrated by overlaying dynamic topography on the geographical
distribution of high positive and negative differences between
climatological values of the Rossby radius. Plainly, there is nothing in the
Arctic Ocean or Nordic seas of comparable strength. In their Fig. S13,
Morison et al. (2012) show Arctic dynamic ocean topography (DOT; akin to dynamic
height), from which it is seen that large-scale (pan-Arctic) horizontal
gradients of DOT are of order 10-7, two orders of magnitude smaller
than Gulf Stream values (e.g. Kelly and Gille, 1990). Near and downstream of
Fram Strait, large horizontal dynamic height gradients, of order 10-6,
exist across the East Greenland Current (EGC; Manley et al., 1992). Since
the PHC search radius depends on the correlation length scale, and is set to
500 km in the Arctic Ocean and 100 km in the Nordic seas, some impact of
horizontal averaging may be seen locally: for example, in the EGC, and also
likely across the Beaufort Gyre (Fig. 4).
Difference between BGEP Rossby radii,
decimated minus full resolution.
The net uncertainty in our PHC-derived estimates of the annual mean Rossby
radius depends, therefore, on the major terms, which are (i) secular
variability, (ii) seasonal variability, and (iii) the quasi-random component
resulting from horizontal averaging. From the BGEP case, we estimate (i) as
10 %; from Fig. 2a and b, we estimate (ii) as 10 %, and from C98, we
estimate (iii) as 5 %, for a root-sum-square total of 15 %.
Finally, PHC is relatively low resolution, at 1∘× 1∘ , so we inspect output from the OCCAM model: fields of
R1 for March and August 1992 are shown in Fig. 6. Using the same
regions as defined in Table 1 and Fig. 3, we find that in no region in
OCCAM is R1 more than 1.5 km different from PHC, with the exception of
the Canadian Arctic Archipelago, where OCCAM results are approximately half
those of PHC. This is not a flaw in the model but rather a consequence of
the model's much higher resolution than PHC. The model is able to capture
the narrow channels – with their low values of R1 – which are absent
from PHC. The major difference between model and PHC, given their
quantitative similarity, lies in the detail visible in the model values: the
imprint of bathymetry on R1 stands out, particularly on the Siberian
shelves, but also over the deep ocean ridges.
(a) OCCAM Rossby radius mode 1 (km) for August 1992. (b) OCCAM Rossby radius mode 1 (km) for March 1992.
DiscussionRossby radius and stratification
The wide range of values of Rossby radii throughout the Arctic Ocean are a
result of the interplay between density stratification and water depth, with
the former largely (but not exclusively) controlled by upper-ocean salinity
variability. We illustrate this as follows.
We consider three cases to illustrate high, medium and low values of
R1. We approximate Eqs. (2) and (6) by setting g= 10 m s-2, f=1.4×10-4 s-1, β=8×10-4 kg m-3 psu-1.
The impact of temperature on density is neglected, since stratification in
and around the Arctic is dominated by salinity (Carmack, 2007). For each
case, we choose representative values of salinity and of scale depth, to
estimate dρ/dz.
High values of R1 – of order 10 km – are found in the deep basins of
the Arctic Ocean. A typical upper-ocean salinity of 32, a deep salinity of
34.8 (e.g. Carmack, 2000) and a scale depth of 1000 m result in dρ/dz∼2×10-3 kg m-4, N∼5×10-3 s-1, and R1∼11 km. Equation (5) helps to
understand the choice of scale depth: the density stratification is very
weak below ∼1000 m, and the vertical integral of N is thus
dominated by the stratification above 1000 m.
Medium values of R1 – of order 5 km – are seen in the central
Greenland Sea (Karstensen et al., 2005). With a surface-to-bottom
(potential) density difference of ∼0.1 kg m-3 and a
scale depth of 3000 m, we find dρ/dz∼3×10-5 kg m-4, N∼0.6×10-3 s-1 and R1∼4 km. This very low value in a deep ocean region is the result of the weak
stratification that pertains throughout the water column. Several
publications have described “sub-mesoscale convective vortices” found in
the Greenland Sea (e.g. Gascard et al., 2002; Wadhams et al., 2002; Budeus et
al., 2004), and these features have radii ca. 5 km. It appears that these
are not, in fact sub-mesoscale but mesoscale. It so happens that the
mesoscale in this basin is very small.
Low values of R1, of order 1 km, are illustrated by considering the East
Siberian Sea (Münchow et al., 1999). The scale depth is set to 50 m and
the surface-to-bottom salinity difference to 2, resulting in
dρ/dz∼3×10-2 kg m-4, N∼2×10-2 s-1 and R1∼2 km. While the density gradient (and
therefore N) is high, the small resulting value of R1 is due to the small
value of “full ocean depth” – around 50 m. With homogenization of the
water column (directly or indirectly) through winter heat loss, the vertical
density gradient can assume very small values (significantly less than 1 km),
which presents a challenge to observations and models alike, given the
importance of the shelf seas to Arctic freshwater fluxes and water mass
structure, and thereby to local and non-local climate.
Difference between exact solution and WKBJ/LG approximate solution
(exact minus WKB) for mode 1 Rossby radius (km), annual mean.
We assess the usefulness of the WKBJ/LG approximate solutions by plotting
the difference field (exact solution minus approximation) for annual mean
values of R1 in Fig. 7. It is seen that the WKBJ/LG method is in error
typically, and over most of the region, by ±1–2 km. This represents
an uncertainty of ∼20 % over the deep basins of the Arctic
Ocean, but is a larger relative uncertainty where R1 is small – in the
Greenland Sea and the shallow shelf seas. Noting that c in the Arctic is
nearly everywhere ca. 1 m s-1 (to within a factor 2; not shown), and using
the above estimates of N, the vertical length scale c/N is 2000 m in the
Greenland Sea and 50 m in the shelf seas. In both cases, this is comparable
to the water depth, so the WKBJ/LG scale assumption is not well satisfied.
There is nothing in the published literature with which to compare our
results. However, Saenko (2006) describes Rossby radii calculated from an
ensemble of coarse (∼1∘ by 1∘)
resolution climate models up to 85∘ N using the WKBJ/LG
approximation, presented as zonal means. This is a highly unsatisfactory
metric in high northern latitudes because it conflates extensive shallow
shelf seas and deep ocean basins. Nevertheless, we parallel this style of
presentation in Fig. 8, which shows annual, summer and winter zonal means
of R1 and R2. Seasonality has little impact by this metric. Minimum
values occur around 65–70∘ N; this latitude band mainly
comprises the southern Nordic seas and Davis Strait, with some elements of
shelf seas. Maxima are found near the Pole (85–90∘ N), and
these waters are all of the deep Arctic Ocean. The results of Saenko (2006)
bear some similarities to this. Models with the smallest mean Rossby radii
are in agreement with ours, but several others show substantial
latitudinally dependent overestimates. Nevertheless, it is encouraging that
some models appear to be capable of producing sensible density
stratification (in the zonal mean).
Observed eddies
There have in the past been several high-resolution surveys of Arctic Ocean
eddies, reported in Newton et al. (1974), Hunkins (1974), Manley and Hunkins (1985),
D'Asaro (1988), Padman et al. (1990), Muench et al. (2000), Pickart
et al. (2005), Timmermans et al. (2008), Nishino et al. (2011) and Kawaguchi
et al. (2012), and stemming (variously) from field programmes such as the
Arctic Ice Dynamics Joint Experiment (AIDJEX) in the 1970s, the Arctic
Internal Wave Experiment (AIWEX) in 1985, Scientific Ice Expedition (SCICEX)
measurements from the 1990s, the Western Arctic Shelf-Basin Interaction
(SBI) programme of 2005, ice-tethered profilers (ITPs), and an expedition on
the R/V Mirai in 2010. Curiously, all these papers report on eddies observed in
the Canada Basin. It is not clear whether the Canada Basin is “infested”
with eddies or whether there is simply a paucity of eddy-resolving
measurements in the other basins, caused by the difficulty of making such
measurements given the ice cover.
Zonal means of PHC Rossby radii modes 1 and 2.
Still, these cited observations are all more or less consistent in how they
describe the observations. The eddy has a core where rotation is (close to)
solid body, and the outer edge of the core defines the radius of maximum
velocity. Further outwards from the edge of the core is a region which is
still rotating but where the velocity progressively reduces (the
“penumbra”; Hoskins et al., 1985), out to a maximum radius of influence.
Typical core radii are ∼7 km, and the typical maximum radius
of influence is ∼15 km. The eddy described by Kawaguchi et al. (2012) was (apparently) an unusually large one, doubling these values.
An empirical quantification of this description is given by Timmermans et al. (2008).
Hoskins et al. (1985) suggest an approach (the pursuit of which is beyond
the scope of the present study) whereby an explanation of these observations
may be developed. At its point of generation, an eddy is in solid-body
rotation. If, for example, the generation mechanism is baroclinic
instability, the Rossby radius would then describe the solid-body rotation
radius because (as noted in Sect. 1) it would characterize the scale of
the waves that grow most rapidly as a result of baroclinic instability.
Subsequently, the closed-contour potential vorticity anomaly induces flow in
the surrounding volume (the penumbra). By inference, therefore, most of the
observed Canadian Basin eddies should be of Mode 2, since the local value of
R2 is ∼6 km (Fig. 2c), similar to the solid-body core
radii, with the exception of the large eddy (Kawaguchi et al., 2012) whose
core radius is similar to the local value of R1. As a further
complication, Chelton et al. (2011) note that eddies may be 2 or 3
times larger than the Rossby radius. It is not straightforward to associate
calculated Rossby radii with observations of eddies.
Final remarks
Timmermans et al. (2008) demonstrate the feasibility of making
quasi-Lagrangian observations of Arctic Ocean eddies with ice-tethered
profilers, but Eulerian measurements present a challenge. The only sustained
Arctic Ocean measurement programme to resolve successfully the local Rossby
radius is located north of Alaska (Nikopoulos et al., 2009) with a typical
moored instrument spacing of ∼5 km. Furthermore, the
logistical and operational constraints of trans-polar hydrographic sections
conducted on research icebreakers mean that they cannot get close to
resolving the Rossby radius (e.g. Carmack et al., 1997).
The main aim of this paper was to present fields of the Rossby radii in the
Arctic Ocean and adjacent seas. Lacking a quantitative appreciation of
Rossby radii, it is possible for features to be “missed” by measurement
programmes. For example, the Shelf Break Branch of the Arctic Circumpolar
Boundary Current has only recently been described (Aksenov et al., 2011).
This is a shallow feature transporting halocline waters that circuits most
of the Arctic Ocean. Over the shelf break the Rossby radius is typically
∼7 km, so this current is sufficiently narrow that it had
slipped almost unnoticed between more widely spaced standard measurement
locations. The model study inspired reanalysis of past measurements, and
deliberate targeting of new measurements. Advances in understanding of
Arctic Ocean circulation and dynamics will likely be found from measurements
and models in combination.
Acknowledgements
This study was funded by the UK Natural Environment Research Council, and is
a contribution to to the UK TEA-COSI project. The PHC data were downloaded
from http://www.psc.apl.washington.edu/Climatology.html, in version 3.0.
BGEP data were downloaded from the project website,
http://www.whoi.edu/beaufortgyre/. The authors are grateful to the editor
and the reviewers for their patience, and to Takamasa Tsubouchi for help
with data wrangling. Calculations were performed and plotted with the SciPy
and matplotlib open source python packages (http://www.scipy.org; http://www.matplotlib.org).
Edited by: M. Hecht
References
Aksenov, Y., Bacon, S., Coward, A. C., and Nurser, A. J. G.: The North
Atlantic Inflow into the Nordic Seas and Arctic Ocean: a high-resolution
model study, J. Marine Sys., 79, 1–22, 2010a.Aksenov, Y., Bacon, S., Coward, A. C., and Holliday, N. P.: Polar outflow
from the Arctic Ocean: a high-resolution model study, J. Marine Sys., 83,
14–37, 10.1016/j.jmarsys.2010.06.007, 2010b.Aksenov, Y., Ivanov, V. V., Nurser, A. J. G., Bacon, S., Polyakov, I. V.,
Coward, A. C., Naveira Garabato, A. C., and Beszczynska-Moeller, A.: The
Arctic Circumpolar Boundary Current, J. Geophys. Res., 116, C09017,
10.1029/2010JC006637, 2011.Budeus, G., Cisewski, B., Ronski, S., Dietrich, D., and Weitere, M.:
Structure and effects of a long lived vortex in the Greenland Sea, Geophys.
Res. Lett., 31, L05304, 10.1029/2003GL017983, 2004.
Carmack, E. C.: The freshwater budget of the Arctic ocean: sources, storage
and sinks, edited by: Lewis, E. L., NATO Adv. Res. Ser., 91–126, 2000.
Carmack, E. C.: The alpha/beta ocean distinction: a perspective on
freshwater fluxes, convection, nutrients and productivity in high-latitude
seas, Deep-Sea Res. II, 54, 2578–2598, 2007.
Carmack, E. C., Aagaard, K., Swift, J. H., Macdonald, R. W., McLaughlin, F.
A., Jones, E. P., Perkin, R. G., Smith, J. N., Ellis, K. M., and Killius, L.
R.: Changes in temperature and tracer distributions within the Arctic Ocean:
results from the 1994 Arctic Ocean section, Deep-Sea Res. II, 44, 1487–1502,
1997.
Chelton, D. B., de Szoeke, R. A., Schlax, M. G., El Naggar, K., and Siwertz,
N.: Geographical variability of the first baroclinic Rossby radius of
deformation, J. Phys. Oceanogr., 28, 433–459, 1998.
Chelton, D. B., Schlax, M. G., and Samelson, R. M.: Global observations of
nonlinear mesoscle eddies, Prog. Oceanogr., 91, 167–216, 2011.
D'Asaro, E. A.: Observations of small eddies in the Beaufort Sea, J.
Geophys. Res., 93, 6669–6684, 1988.
Gascard, J.-C., Watson, A. J., Messias, M.-J., Olsson, K. A., Johannessen,
J., and Simonsen, K.: Long-lived vortices as a mode of deep ventilation in
the Greenland Sea, Nature, 416, 525–527, 2002.
Giles, K. A., Laxon, S. W., Ridout, A. L., Wingham, D. J., and Bacon, S.:
Western Arctic Ocean freshwater storage increased by wind-driven spin-up of
the Beaufort Gyre, Nat. Geosci., 5, 194–197, 2012.
Gill, A. E.: Atmosphere-Ocean Dynamics, Academic Press, 662 pp., 1982.
Hallberg, R.: Using a resolution function to regulate parameterizations of
oceanic mesoscale eddy effects, Ocean Model., 72, 92–103, 2013.
Hecht, M. W. and Smith, R. D.: Toward a physical understanding of the North
Atlantic: a review of model studies in an eddying regime, in:
Ocean Modeling in an Eddying Regime, edited by: Hecht, M. W. and Hasumi, H.,
Geophys. Monog. Series, 177, 231–239, 2008.
Hoskins, B. J., McIntyre, M. E., and Robertson, A. W.: On the use and
significance of isentropic potential vorticity maps, Q. J. Roy. Meteor. Soc., 111, 877–946, 1985.
Hunkins, K. L.: Subsurface eddies in the Arctic Ocean, Deep-Sea Res., 21,
1017–1033, 1974.
Jakobsson, M., Cherkis, N., Woodward, J., Macnab, R., and Coakley, B.: New
grid of Arctic bathymetry aids scientists and mapmakers, Eos Trans. AGU, 81, 89–96, 2000.Karstensen, J., Schlosser, P., Wallace, D. W. R., Bullister, J. L., and
Blindheim, J.: Water mass transformation in the Greenland Sea during the
1990s, J. Geophys. Res., 110, C07022, 10.1029/2004JC002510, 2005.
Kawaguchi, Y., Itoh, M., and Nishino, S.: Detailed survey of a large
baroclinic eddy with extremely high temperatures in the western Canada
Basin, Deep-Sea Res. I, 66, 90–102, 2012.Kelly, K. A. and Gille, S. T.: Gulf Stream surface transport and statistics
at 69∘ W from the Geosat altimeter, J. Geophys. Res., 95, 3149–3161, 1990.
Manley, T. O. and Hunkins, K.: Mesoscale eddies of the Arctic Ocean, J.
Geophys. Res., 90, 4911–4930, 1985.
Manley, T. O., Bourke, R. H., and Hunkins, K. L.: Near-surface circulation
over the Yermak Plateau in northern Fram Strait, J. Marine Systems, 3, 107–125, 1992.
Marsh, R., de Cuevas, B. A., Coward, A. C., Jacquin, J., J. Hirschi, J.-M.,
Aksenov, Y., Nurser, A. J. G., and Josey, S. A.: Recent changes in the North
Atlantic circulation simulated with eddy- permitting and eddy-resolving
ocean models, Ocean Model., 28, 226–239, 2009.
Morison, J., Kwok, R., Peralta-Ferriz, C., Alkire, M., Rigor, I., Andersen,
R., and Steele, M.: Changing Arctic Ocean freshwater pathways, Nature, 481,
66–70, 2012.
Muench, R. D., Gunn, J. T., Whitledge, T. E., Schlosser, P., and Smethie
Jr., W.: An Arctic cold core eddy, J. Geophys. Res., 105, 23997–24006, 2000.
Münchow, A., Weingartner, T. J., and Cooper, L. W.: The summer
hydrography and surface circulation of the East Siberian Shelf Sea, J. Phys.
Oceanogr., 29, 2167–2182, 1999.
Newton, J. L., Aagaard, K., and Coachman, L. K.: Baroclinic eddies in the
Arctic Ocean, Deep-Sea Res., 21, 707–719, 1974.Nikopoulos, A., Pickart, R. S., Fratantoni, P. S., Shimada, K., Torres, D.
J., and Jones, E. P.: The western Arctic boundary current at 152∘ W:
structure, variability and transport, Deep-Sea Res. II, 56, 1164–1181,
10.1016/j.dsr2.2008.10.014, 2009.Nishino, S., Itoh, M., Kawaguchi, Y., Kikuchi, T., and Aoyama, M.: Impact of
an unusually large warm-core eddy on distributions of nutrients and
phytoplankton in the southwestern Canada Basin during late summer/early
fall 2010, Geophys. Res. Lett., 38, L16602, 10.1029/2011GL047885, 2011.
Padman, L., Levine, M., Dillon, T., Morison, J., and Pinkel, R.: Hydrography
and microstructure of an Arctic cyclonic eddy, J. Geophys. Res., 95,
9411–9420, 1990.Pickart, R. S., Weingartner, T. J., Pratt, L. J., Zimmermann, S., and
Torres, D. J.: Flow of winter-transformed Pacific water into the western
Arctic, Deep-Sea Res. II, 52, 3175–3198, 2005.
Proshutinsky, A., Krishfield, R., Timmermans, M.-L., Toole, J., Carmack, E.,
McLaughlin, F., Williams, W. J., Zimmermann, S., Itoh, M., and Shimada, K.:
Beaufort Gyre freshwater reservoir: state and variability from observations,
J. Geophys. Res., 114, C00A10, 10.1029/2008JC005104, 2009.
Rabe, B., Karcher, M., Schauer, U., Toole, J. M., Krishfield, R. A.,
Pisarev, S., Kauker, F., Gerdes, R., and Kikuchi, T.: An assessment of
Arctic Ocean freshwater content changes from the 1990s to the 2006–2008
period, Deep-Sea Res. I, 58, 173–185, 2011.
Saenko, O. A.: Influence of global warming on baroclinic Rossby radius in
the ocean: a model intercomparison, J. Climate, 19, 1354–1360, 2006.Smith, R. D., Maltrud, M. E., Bryan, F. O., and Hecht, M. W.: Numerical
simulation of the North Atlantic Ocean at 1/10∘, J. Phys.
Oceanogr., 30, 1532–1561, 2000.
Smith, W. H. F. and Sandwell, D. T.: Global seafloor topography from
satellite altimetry and ship depth soundings, Science, 277, 1957–1962, 1997.
Steele, M., Morley, R., and Ermold, W.: PHC: A global ocean hydrography with
a high quality Arctic Ocean, J. Climate, 14, 2079–2087, 2001.
Stroeve, J., Serreze, M., Drobot, S., Gearheard, S., Holland, M., Maslanik,
J., Meier, W., and Scambos, T.: Arctic sea ice extent plummets in 2007, Eos,
89, 13–20, 2008.
Timmermans, M.-L., Toole, J., Proshutinsky, A., Krishfield, R., and
Plueddemann, A.: Eddies in the Canada Basin, Arctic Ocean, observed from
Ice-Tethered Profilers, J. Phys. Oceanogr., 38, 133–145, 2008.Wadhams, P., Holfort, J., Hansen, E., and Wilkinson, J. P.: A deep
convective chimney in the winter Greenland Sea, Geophys. Res. Lett., 29, 1434,
10.1029/2001GL014306, 2002.