The Interaction of an Ocean Eddy With an Ice Edge Ocean Jet in a

JOURNAL OF GEOPHYSICAL
RESEARCH, VOL. 96, NO. C3, PAGES 4675-4689, MARCH 15, 1991
The Interaction of an Ocean Eddy With an Ice Edge Ocean Jet
in a Marginal Ice Zone
DAVID
C. SMITH
IV AND ARLENE A. BIRD
Departmentof Oceanography,
Naval PostgraduateSchool,Monterey,California
Remotelysensedimagesin the EastGreenlandCurrentmarginalice zoneregionindicatea variety
of mesoscalemotions.These rangefrom isolatedmonopolesand dipolesto multiple eddy fields.
Recentobservationsindicatethat many of the eddiesmay be advectedinto the regionfrom the West
Spitsbergen
Currentregion.TheytheninteractwiththeEastGreenland
Currentmarginalicezoneand
its associated
along-ice-edge
jet. In thispaperwe illustratehowtheice floefieldscanevolveduringthe
interactionof an oceaneddy with an along-ice-edge
oceanjet. The effectsof eddy senseof rotation,
flowstrength,andbottomtopography
areconsidered
in barotropicandbaroclinic,two-layernumerical
experiments.
Initially, barotropicexperiments
involvethe combinedeffectsof flat-bottomeddy-jet
interactionsand eddy-topography
interactions.Over a flat bottom, anticyclonesinteract with the
cyclonicallyshearedsideof the jet and moveseaward,exportingice awayfrom the ice edgein a
dipolardistribution.Cyclonesmovetowardthejet andare advecteddownstream.
A seconddipole
formation mechanismis associatedwith eddy-topographyinteractionand is a consequenceof the
shallowpycnoclinedepthof the region.Rapidtopographic
Rossbywave radiationin the lower layer
spreadsenergyalongslopeand upslope(downslope)for cyclones(anticyclones).
The upperocean
response
to thisprocessis a dipolewhichpropagates
parallelto the ice edgewith shallowwaterto the
left (topographic
eastward).Jet-induced
andtopography-induced
eddypropagation
tendencies
combine to enhancecyclone-jetinteractionand limit anticyclone-jet
interactions.The dipoleformation
associated
witheddy-topography
interactionis weakenedby thepresenceof thejet but noteliminated.
Eddy-induced
distortions
in the ice edgeare suppressed
by the stabilizing
effectof topography
on the
jet. Sensitivityof the initiallybarotropicresultsto variationin the Rossbynumberand degreeof
topographicslopeillustratesthat eddiesare more efficientin producingice edgedistortionsfor
strongerflowsand weakertopographic
slopes.For initiallybaroclinicsimulations
over topography,
resultsare comparableto flat-bottomcases.Strongupper layer eddy-jet interactionscan occur
independentof topographiceffectswhich are constrainedto the lower layer.
1.
ties of the front or the broader
INTRODUCTION
An overview of various types of mesoscalemotions seen
in remotely sensedimagesof the marginalice zone (MIZ) is
givenby Ginsburgand Federov [1989].The motionsinclude
monopole, vortex dipole, and mesoscaleocean jets transverse to the ice edge. The spatial scalesof the eddies can
range from several kilometers to 100 km. Ginsburg and
Federov [1989] suggestthat the eddies are short-livedwith
lifetimes of only several days. Many of the eddies they
discuss are dipoles (see, for example, Figure 1). They
reinterpretthe earlier observationof cyclonesby Wadhams
and Squire [1983] and Johannessenet al. [1983] as being
dipolesin which only the cyclonicportion of the dipole is
tracedby the ice. The formationof dipolesthey speculateis
not likely related to baroclinic or barotropic instabilitiesor
bottom topography.Other processessuch as near-surface
density front instability [Griffiths and Linden, 1981], upwelling along an irregular ice edge [Hiikkinen, 1986], or
uneven ice edge melting (with subsequentgeostrophicadjustment of the melted pool) may be related to the dipole
formations.Dipole formationsare also seenin other regions
in satellite images. In the Oyashio region, Federov and
Ginsberg [1989] speculatethat dipole formations may be
related to eddy interaction with the Oyashio.
The existence of open ocean eddies adjacent to the ice
edge front was documentedby Johannessenet al. [1983,
1987]. In those studiesit was argued that dynamic instabili-
East Greenland
Current
(EGC) were responsiblefor the existence of the eddies.
Recently, Gascardet al. [1988]andFoldvik et al. [1989]have
suggestedthat the eddiesare not locally producedbut are
advected into the East Greenland Current region across
Fram Strait from the West SpitsbergenCurrent region. This
suggests
that the eddiescouldinteractwith the EGC and the
ice edgefront. Motivatedby the suggestions
of Federovand
Ginsberg[1989],and the driftingbuoy resultsof Gascardet
al. [1989], we investigatethe interactionof an existingeddy
with an along-ice-edgeoceanjet. We considerthe effectsof
eddyandjet strength,eddy senseof rotation,initial vertical
shear, and the presenceof a topographicslope on eddy-jet
interactions.The signatureof these processesin the ice
concentration field is illustrated and the physical mechanismsresponsibleare examined.
Ikeda and Lygre [1989]alsostudiededdy-jetinteractions,
but in the context of a coastaljet flowing over a topographic
sill. Their jet thus flows acrossisobathsinto deeper water,
while ours flows alongisobaths.Emphasisin their study is
on eddy interactions with meanders of a baroclinically
unstablejet. They showhow eddiescan modifythe meander
wavelengthof the jet. The interactionsreportedhere apply
to the periodbeforeunstablejet meanderingoccurs.Emphasis is on eddy motion and the resultantsignaturein the ice
edge.
•
the American Geophysical Union.
In section 2 the model and experiment initial conditions
are described. Flat-bottom eddy-jet experiments are discussed in section 3. The effect of topography on eddy
evolution is consideredin section 4. Eddy-jet interactions
Paper number 90JC02262.
over a slopingbottomare discussedin section5. In section
This paper is not subjectto U.S. copyright.Publishedin 1991by
4675
4676
SMITH AND BIRD: EDDY-JET
78o30.N
INTERACTION
IN A MARGINAL
ICE ZONE
.
MIZEX
87
1 April1987
1630 - 1900 UT
--.
77"00'N
Fig. 1. Dipole formation as seen in a synthetic aperture radar image [from Shuchman et al., 1988].
6, the sensitivity of the results to flow strength and degree of
topographic slope is discussed.Experiments in sections3, 4,
5, and 6 are for initially barotropic flow fields. Baroclinic
experiments are discussedin section 7, and conclusionsare
provided in section 8.
2.
2.1.
NUMERICAL TECHNIQUE AND MODEL PARAMETERS
Ocean
Model
Experiments are performed using a two-layer primitive
equation numerical scheme for the ocean with a coupled
one-layer ice model. The model is essentially the same as
that used by Smith et al. [ 1988] to study the interaction of an
isolated eddy with a marginal ice zone. The reader interested
in model developmentand testingis referred to that study for
further
details.
Motion in each ocean layer is governed by a momentum
equation
at
ot
•7' Vi=O
for layer thickness hi (i = 1, upper layer; i = 2, lower),
transports Vi, and velocities •i- The fluid is hydrostatic,
Boussinesq,and the fluid density Pi in each immisciblelayer
is fixed. The equations are finite differenced in a semiimplicit numerical scheme following Hurlburt and Thompson [1980], who discuss it more fully.
The ocean is coupled to the ice through an ice-water
interfacial stress riw:
?iw=t91Ciw(ttl_
ttw)lttl_ttwl
for ice (Ul) and upper layer ocean (Uw) velocity vectors. The
constantdragcoefficients,(PI, Ciw,etc.) are chosenfollowing
Hiikkinen [1986] and are given in the notation section.
Subgrid-scaledissipationalprocessesare represented by a
horizontal Laplacian operator on transport. The dissipation
coefficient
Ah is 10m2 s-1. A dissipation
timescaleassoci-
ovi
•+
Ohi
•+
(V' V i+ V i' •7)• i + k X fV i
ated with A h is
L 2/gh -- 30 days
= -hiVPi + AhV2Vi+ • + A?iw
Pi
and a continuity equation
All notation
is defined
in the notation
section.
A rectangular (110 x 80 km) finite difference gridded
domain is used. Grid spatial resolution (2Ax) is 2 km. The
SMITH AND BIRD: EDDY-JET
TABLE
1.
Experiment Initial Parameters
Rotation
cm s
-1
Ro
ICE ZONE
4677
barotropic (v• = 1,,2)velocity distributions. The barotropic
eddy and jet are defined as
Topographic
•'max,
Experiment
INTERACTION IN A MARGINAL
Slope/5 hi(y)= I•l
Flat Bottom, Barotropic, Eddy-Jet (Section 3)
ACJ
CJ
ACJ2
CJ2
ACJ3
CJ3
AC
C
AC
C
AC
C
25/25
25/25
12/12
12/12
37/37
37/37
0.34
0.34
0.17
0.17
0.51
0.51
0
0
0
0
0
0
2Lj2
Y>Y0
BF2
Broad TopographicSlope, Barotropic Eddy (Section4)
ACB
AC
25/25
0.34
1.5 x 10-5
CB
C
25/25
0.34
1.5 x 10-5
where
r2 = (x - Xc)2 + (y - yc)2
Broad Topographic Slope, Barotropic Eddy-Jet (Sections 5, 6)
ACJB
AC
25/25
0.34
1.5 x 10-5
CJB
C
25/25
0.34
1.5 x 10-5
ACJB2
AC
12/12
0.17
1.5 x 10-5
CJB2
C
12/12
0.17
1.5 x 10-5
ACJB3
AC
37/37
0.51
1.5 x 10-5
CJB3
C
37/37
0.51
1.5 x 10-5
(negative for cyclones and positive for anticyclones) for
layer thicknessvalueshi (i = 1, 2) andXc, Ycare coordinates
of the eddy center. Although in sections3 through 6 the eddy
and jet are initialized with barotropic structure, baroclinic
structure associatedwith the layer interface is free to evolve.
The seawardedge of the jet is at Y0 (Y grid pointj = 38, 38
km from the lower boundary of the model domain; see
Figure 2). The eddy center is initially 9 km from the seaward
Narrow Topographic Slope, Barotropic Eddy-Jet (Section 6)
ACJN
AC
25/25
0.34
3.0 x 10-5
CJN
C
25/25
0.34
3.0 x 10-5
ACJN2
AC
12/12
0.17
3.0 x 10-5
CJN2
C
12/12
0.17
3.0 x 10-5
ACJN3
AC
37/37
0.51
3.0 x 10-5
CJN3
C
37/37
0.51
3.0 x 10-5
edgeof thejet. Lj andL e arethe e-foldingwidthscales(=5
km) for the jet and eddy, respectively. The amplitude of the
Gaussian distribution was chosen to give a maximum veloc-
Broad Topographic Slope, Baroclinic Eddy-Jet (Section 7)
BCACJB
AC
40/10
0.56/0.14
1.5 x 10-5
BCCJB
C
40/10
0.56/0.14
1.5 x 10-5
ity (I2max)
of approximately
25cms-1 fortheeddyandjet in
mostexperiments
(sections3-5). Weaker(12 cm s-1) and
stronger
(37cms-1) casesareshownin section6. Theinitial
AC, anticyclonic'C, cyclonic, /5= (1/H2)(Oh2/Oy).
velocity and eddy sense of rotation for each experiment is
given in Table 1.
A radiation
condition
was used on both the downstream
(left) and upstream (right) boundaries [Camerlengo and
O'Brien, 1980]. The zonal boundaries are no-slip walls where
estimatedby 13L2/U[Flied et al., 1983].For the flow both tangential and normal flow are set equal to zero.
Consistent with the short lifetimes of the dipoles and
parametersin these experiments, this quantity is of the order
of 10-3 indicating
that the effectof /3 is negligible.
The transversejets suggestedby Ginsberg and Federov [1989],
Coriolis parameter is thus a constant in this study.
the simulationswere integratedfor short periods (<5 days).
Table 1 provides initial conditions for the experiments. The thermocline depth in the East Greenland Current ranges
The initial state consists of a geostrophically balanced jet from 25 to 100 m. Here the upper layer mean thickness is
extending uniformly across the basin aligned with the ice chosen to be 50 m. The lower layer mean thickness is 4000
edge (defined below). A Gaussian open ocean eddy in m. The first internal Rossby radius of deformation(Re) is
approximately 7 km for these layer thickness choices.
gradient balance is initially seaward of the ice edge jet.
Experiments in sections 3, 4, 5, and 6 are initialized with
According to Gascard et al. [1988], the East Greenland
domain is centered at latitude 80øN. The importance of
planetary /3 in a quasi-geostrophicvorticity equation, is
free-drift
25
MIZ
50
cm/s
m
ocean
jet
4000
sloping
80
km
Fig. 2.
•
•
m
bottom
110 km
Schematicillustration of the coupled ice-oceanmodel and initial condition for the experiments.
4678
SMITH AND BIRD: EDDY-JET INTERACTION IN A MARGINAL ICE ZONE
Current has a narrow jetlike structurewith width of about 25
km, flowing over the continental slopebetween the 1000and
2000 m isobaths. Our dimensional jet width of 5 km is
narrower than that cited by Gascard et al. [1988]; however,
a nondimensional
jet widthLj/Rdfor ourjet isof theorderof
1, as Gascard et al. [ 1988]indicate is appropriatefor the East
Greenland Current jet.
mid-domain (see Figure 3a, day 0). The ice is initialized with
the ocean velocity. Wind-forcing effects are not considered
in this study. The ice acts largely as a passivetracer for the
ocean motion. Slight differences between the ice and ocean
fields may be expected due to differences between dissipation coefficients
3.
2.2.
in the ice and ocean.
FLAT BOTTOM, BAROTROPICEDDY-JET
INTERACTIONS
Ice Model
The interaction of an eddy with an adjacent oceanjet was
consideredby Stern and Flierl [ 1987]. In that study, an eddy
representedby a point vortex was allowed to interact with a
c)tlI
c)tlI
c)tlI
A .
• + Ul
+ Vl
= fvl - --r• w
jet representedby a frontal boundary between two fluids of
Ot •x
•
m
uniform but different potential vorticity. Their study illustrated that an anticyclone close to the cyclonically sheared
O(hl + h2 + d)
-- g
+ AhV2Ui side of a jet can pair with the cyclonically shearedfluid of the
Ox
jet and propagatelaterally away from the axis of the jet. In
contrast, a cyclone on the cyclonically sheared side of a jet
01•I
c)1•
I
01•I
A .
moves toward the jet and is advected downstream with the
+tll
+ Pl
= _fll I __ ,w
jet. Becausethe experimentswere done on anf plane with a
at
•
-•y
mry
quiescentbackgroundflow field, the jet-induced eddy propO(hl + h2 + d)
agation represents an eddy propagation mechanism not
-- #
+ AhV2•i associatedwith planetary rotation or advection by broader
oy
mean flows. These results were extended by Smith and
and continuity equations
Davis [1989] to show that these processeswere also seen
when eddies and jets with more realistic finite lateral vortica A O(A ui) O(A•,i)
ity
distributions were considered.
•+•+•
= AaV2A
at
ox
Oy
Stern and Flierl [1987] showed that significanteddy motion away from a jet can occur for eddies of opposite sign
am O(mui) O(m•,i)
vorticity of the jet edge when the eddy-jet separationdis• +•
+•
= AmV2m
tance is less than a length scale l, defined as the squareroot
at
ox
Oy
of the vortex circulation divided by the jet shear potential
for ice concentration A and ice mass m. As in previous ice vorticity. Stern and Flierl [1987] showed that this is equivmodel studies [Hdikkinen, 1987; Hibler, 1979], a Laplacian alentto • r fordimensional
eddyradius
r. In thisstudy,r
damping term has been included in the continuity equations is 5 km, making l equivalent to 9 km. For barotropic point
Motion in the ice is governed by the momentumequations
for
a and
m.
The
Peclet
number
indicates
the
relative
importance of advection to diffusionin the ice concentration
field. For the values chosen here, we estimate
UL
(0.2)(5000)
Aa
(30)
Pe ....
30
In other studies of tracer evolution over mesoscale motions,
a Peclet number > 200 is considered to indicate nondiffusive
motions [Mied et al., 1990]. Our ice concentration field
cannot be considered
nondiffusive.
The ice equationscontain a pressureforce associatedwith
sea surface slope. Most mesoscalemodeling studiesof the
MIZ have neglected this term. Its importance in no-wind
simulationsis demonstratedby Smith et al. [1988]. Although
no internal ice stress [ttibler, 1979, 1984] is included in the
ice equations, the effects of viscous-plasticice rheology are
being considered in a separate study (W. P. Budgell et al.,
manuscript in preparation, 1990). Thermodynamic effects
are not consideredin this study. The ice thicknessdistribution D is initially specifiedto be 2 m thick but is then allowed
to vary according to
m
(PI A)
Ice concentration is initially a linear function of the y
coordinate varying from 0.25 to 0.75 in a 10-km band in
vortices, Stern and Flierl show that eddy motion is normal
away from the jet when a nondimensional eddy-jet separation distance R(0) is less than 1.25. R(0) is defined as the
dimensional, initial eddy-jet separation distance, L(0), normalized by I defined above. Smith and Davis [1989] found
that weaker eddy motions away from the jet can occur for
R(0) up to 2.4, beyond which the jet-induced eddy motions
become substantiallyweaker. For the eddy dimension(5-km
radius) chosenhere, this correspondsto a dimensionaleddy
jet separationL(0) distance of 21 km. Eddies further than
this distancewould not experiencejet-induced eddy motion.
In the East Greenland Current MIZ region, eddies on the
seawardsideof the ice edgejet would encounterthe cyclonic
shear of the jet. From these considerations, we anticipate
that anticyclones will acquire seaward motion, while cyclones should move toward the jet and be advected downstream with the jet.
3.1.
Barotropic Anticyclone-Jet Interactions
Figure 3 showsan anticyclone-jet experiment (experiment
ACJ, Table 1). Following Stern and Flierl [1987] and Smith
and Davis [1989], the anticyclone pairs with the cyclonic
vorticity (Figure 3c) of the jet and moves away from the jet.
Figure 4 shows the trajectory for the eddy center for this
experiment. Correspondingaverage translation speedsare
listed in Table 2. L(0) for the experiment in Figure 3 is 9 km,
SMITH AND BIRD: EDDY-JET INTERACTION IN A MARGINAL ICE ZONE
4679
0.75
0.30
_--;;7-7-_-_-_--z-_zzz
-
.........
.........
O.
•o.6•0
2.75
--
1. oo
day 2
day 0
day 4
Fig. 3. Barotropic
flat-bottom
anticyclone-jet
experiment
ACJ, 1/ma
x = 25 cm s-1, days0, 2, and4. (a) Ice
concentration(contourinterval CI = 0.05); (b) upper layer relative vorticity (dashedand solidcontoursare anticyclonic
andcyclonic,
respectively;
CI = 0.1 x 10-4 s-l); (c) surface
heightanomaly
(CI = 0.25cm).
makingR(0) = 1. The motion of the eddy away from the jet
is thus consistent with the previous results of Stern and
Flierl [1987]. In the fiat-bottom barotropic experiments the
fluid evolves barotropically, with motion in each layer
remaining in phase (Figures 3b and 3c).
The interaction of the eddy with the jet also produces an
azimuthal mode one asymmetry in the eddy. The effect of
azimuthal mode one asymmetry on eddy motion can be seen
by decomposingthe eddy into axisymmetric and asymmetric
parts. The axisymmetric part on an f plane has no eddy
propagation tendency. The asymmetric part (azimuthal
mode 1) is dipolaf and nonlinearly advects the symmetric
portionof theeddy.Figure5a shows
theanticyclone
at day
4 of experiment ACJ centered in a box which is 20 km on a
side; Figure 5b shows the eddy after the radially symmetric
part (mode 0) has been removed. The distribution of this
asymmetriccomponentis predominantly dipolaf, confirming
the presenceof azimuthalmode one distortion.The direction
of eddy propagationis along the axis of the dipole pair, and
this augments the jet-induced motion of the anticyclone
awayfromthe cyclonicsideof thejet. Effectsof azimuthal
mode 1 distortions in inducing eddy motion are further
examinedby SmithandBird [1989].SmithandDavis [ 1989]
60
showed that the eddies perturbed in eddy-jet interactions
adjust to axisymmetric once they propagate away from the
jet.
The anticyclone-jet interaction creates an upper ocean
'• 5O
•
CJB
40
TABLE 2.
•
o
CJ
•%
3o
Propagation
Speed,kmd-1
•o
lO
50
Average Eddy Propagation Speeds
loo
Along-slope Distance (km)
Fig. 4. Eddy trajectory for flat-bottom (ACJ, CJ) and topographic
(ACJB, CJB) cases. Symbols are marked at daily intervals.
Layer
x Component
y Component
ACJ upper
CJ upper
ACJB upper
CJB upper
CB upper
-0.25
- 1.50
1.00
-6.33
0.74
-2.50
0.89
- 1.73
1.17
2.00
CB lower
- 3.11
ACB upper
ACB
lower
1.13
- 1.76
2.28
- 1.06
-2.32
4680
SMITH AND BIRD' EDDY-JET INTERACTION IN A MARGINAL ICE ZONE
jet and is advecteddownstream(Figure 4) with the ice edge
jet. The inward and downstream advection of the cyclone
with the jet is consistent with the results of Stern and Fiieri
[ 1987]. Downstream of the cyclone, an ice edgeperturbation
ß (a [•',
,,! :,,
[-i
•
•:,',:,.,,_'_-.',•,,-'
o
with cyclonic vorticity extends seaward. It does not how[•
l:', "'
ever have the appearanceof a vortex pair as in the previous
anticyclone-jet experiment. The ice edge perturbation is
advected downstream with the ice edge jet and does not
.......I
-grow seaward with time. Further downstream a growing
Fig. $. Azimuthal decompositionof barotropiceddy ACJ at day
4. (a) Eddy surface height anomaly centered in a 20-km-square meander perturbed by the eddy reaches the downstream
region. (b) Asymmetric componentafter the axisymmetriccompo- boundary by day 4.
nent has been subtracted.
The results of Gascard et ai. [1988] suggest several
processesassociatedwith the cyclone-jet interactions with
which we are able to compare our results. The drifter
trajectories
and satellite images reported there indicate a
"dipole" which advects ice (Figure 3a) seawardaway from
the ice edge. Ice is concentrated only over the cyclonic seriesof seawardextendingice edgetonguesdownstreamof
portion of the dipole. This ice is initially over the cycloni- cyclone-jet interactions. The tongues have an along-ice
cally sheared portion of the jet. The anticyclone is never wavelength of approximately 50 km. Gascard et ai. [1989]
under the ice and does not acquire any ice during the suggest that the occurrence of these ice-edge tongues is
eddy-jet interaction. While the upper ocean flow field ap- attributable to a destabilizationof the ice edge front by the
pears dipolar in this experiment, the rotational velocity in cyclone-jet interaction. Downstream of the cyclone a tongue
the cyclone is very much weaker than in the originalanticy- associatedwith the off-ice eddy flow is seen [see Gascard et
clone. This is true in numerousother experimentsinvolving ai., 1988, Figure 16c]. Gascard et ai. [1988] note that some
the interaction of anticyclones with the cyclonic side of a ice edge eddies appear to remain stationary, presumably
MIZ ice edgejet. For the modelparameterschosenhere, the topographically trapped by the Hovegard and Greenland
o-. • 2.-.
I
;50,
',,
anticyclone-jetinteractionmechanismfor producingdipoles fracture zones. Our simulationsdo not include these trapalways producesanticycloneasymmetry (usingthe terminol- ping effects.
ogy of Federov and Ginsberg [1989] to denote that the
anticycloneis the strongerof the dipolepaiD. It is interesting
4.
EDDY-TOPOGRAPHY
EXPERIMENTS
to note that the advection of ice over the cyclonic portion of
The strongbarotropic componentof the West Spitsbergen
the dipole in Figure 3 gives the appearanceof a cyclone
asymmetry, while the ocean dipole actually has anticyclone Current and its associated eddies is mentioned by many
asymmetry.
authors. Hence, as the eddies are advected into the EGC
Also evident in this simulation is a broadening of the MIZ region, the continental slope of East Greenland may
marginal ice zone with the appearance of a low-concentra- influence the eddy-ice edge front interactions discussedin
tion region. This low-concentrationregion is associatedwith the previous section. To study these effects a linear sloping
differential Ekman drift over the oceanjet. Maximum cross- bottom is includedin subsequentexperiments. Depth ranges
ice-edge Ekman drift occurs over the jet axis leading to a from 1025 m at the top of the domain to 4025 m at the lower
spreading in the ice concentration field. This effect was lateral boundary. The effect of a slopingbottom is to provide
discussedin the study of wind-driven ice edge jets by a strong vorticity gradient absent in the f plane simulations
above. This vorticity gradient can exert a stabilizing effect
Smedstadt and ROed [1985].
on the jet. It also provides a mechanismfor eddy propagation and decay absent in the flat-bottom experiments. The
3.2. Barotropic Cyclone-Jet Interactions
topographic vorticity gradient associated with this slope
While the experiment in Figure 3 illustrates a mechanism •T = -(fo/H2)(ah2/ay)= 2 x 10-9 m-• s-•.
for dipolar export of ice away from the ice edgethe appliPrevious studies of eddy-topography interaction indicate
cability of this result to the East Greenland Current MIZ is that an eddy over a slopingbottom can decay by topographic
unclear. The drifter trajectories of Gascard et al. [1988] Rossby wave radiation. Radiation leads to eddy azimuthal
indicate that eddiesfrom the West Spitsbergenregion are mode one distortions which provide cross-slope eddy moadvected into the EGC MIZ region. The drifters, however, tions. In the context of Loop Current eddies in the Gulf of
indicate that the eddies are predominantly cyclonic. The Mexico, Smith [ 1986] showed that these distortions lead to a
prior observations of Johannessenet ai. [1987] also indicate downslope (upslope) propagationtendency for anticyclones
that the majority of the eddies are cyclonic. This places (cyclones). In that study, this cross-slopepropagation occyclonic eddiesin contact with the cyclonic side of the East curred whenever the lower layer flow strength exceeded a
Greenland ice edgejet.
critical value. Lower layer flow strengthwas quantified by a
The experiment in Figure 3 was repeated with a cyclonic Rossby number Ro2. For Ro2 > 0.07, cross-slopeeddy
eddy in the initial condition. All other parameters, such as motions were found. The lower layer Rossby number for
eddy-jet separationdistanceL(0) and eddy andjet strength, parameters
chosen
here(l•2max
= 0.25cms-1 L = 5 km)is
are the same. Experiments with a cyclone adjacent to the 0.34. Wave radiation also disperses energy along slope,
cyclonic side of the ice edge front do not exhibit seaward leading to a rapid decay of the original eddy [Louis and
dipole formation as in the anticycloneexperimentshownin Smith, 1982]. A dispersivedecay time scale can be estimated
Figure 3. The time evolutionof this experiment(experiment by 1/13rLand is of the order of days for our parameters.In
CJ, Figure 6b) indicatesthat the cyclone moves toward the contrast, the dispersive decay scale for mid-latitude eddies
SMITH AND BIRD: EDDY-JET INTERACTION IN A MARGINAL ICE ZONE
4681
0.30•
2.5--
day0
day2
day4
Fig. 6. Barotropic
flat-bottom
cyclone-jet
experiment
CJ,Vma
x -- 25cms-l , days0, 2, and4. (a)Iceconcentration
(CI = 0.05);(b) upperlayerrelativevorticity(CI = 0.1 x 10-4 s-l); (c) surface
heightanomaly
(CI = 0.5 cm)
(CI--0.1 x 10-4
associatedwith planetary/3 is of the order of months. The
cross-slopetendenciesmentioned above are efficient only
for that period before the lower layer eddy decays(in this
study, several days).
For comparisonwith subsequenteddy-jet-topographyex-
Lower layer relative vorticity plots for an anticyclone (experimentACB, Figure 7b) showthe typical dispersionpattern seen in numerousprevious isolated eddy studies. The
lower layer anticyclonemovesto the "topographic" south-
westat a speed
whichaverages
2.9kmd-• . The.trajectories
periments,we first show the behaviorof isolatedbarotropic for the upperand lower layer vorticity extremaare shownin
eddies over topography without the ice edge jet. These Figure 8. In contrastto the lower layer eddy path, the upper
experimentsillustrateanotherdipoleformationmechanism. layer anticyclonepairs with an upper layer cyclone to its
day2
day3
day4
Fig.7. Anticyclone
topography
(no-jet)
experiment
ACB.(a)Upper
layerrelative
vorticity
(CI = 0.1x 10
(b)lowerlayerrelativevorticity
(CI = 0.1 x 10
4682
SMITH
AND BIRD'
EDDY-JET
INTERACTION
IN A MARGINAL
ICE ZONE
60
fo
2L2 ] Op2
OdOtV2p2+•(pl-P2)
CB•
50-
9'H2
•..
0x
CB•
9'H 2
+•'•J P2,V2p2+•(Pl-P2)
=0
'X
40-
ACB•
30ACB2
20-
10
+•
In flat-bottom, /3 plane experiments, planetary Rossby
waves exist in both layers associated with the /3 term
(t3 (Op/O
x) in dimensionalform).
In our experiments,the planetary/3 term in the lower layer
is replacedby a topographic/3 (+fod/H) term. The lower
layer dispersionis thus related to topographicRossby dispersion. In contrast to those previous experiments, the
upper vortex in our experiments does not disperse, as f is
constantin the upper layer equation. For the layer thickness
choiceshere H l << H2, making
iI
do
do
4o
Along-slope Distance
Fig. 8. Upper and lower layer trajectories for vorticity extrema
in eddy-topography cases ACB and CB. Subscripts on experiment
labels denote upper (1) or lower (2) layer trajectory.
fo2L
2
fo2L
2
9'H 2
9'H1
(=6 x 10-3)<<•
(=0.5)
Thus the vortex stretchingterm has a negligiblerole in lower
layer dynamics, while upper layer motions receive strong
vorticity input from the lower layer by vortex stretching at
the interface.
While these experiments illustrate another dipole formation mechanism which may be appropriate for the East
north and propagates to the "topographic" east (shallow
Greenlandmarginal ice zone, further analysisis necessaryto
wateron the left) at 1.1km d-I (Table2). Velocityvectors determine the robustnessof this result. A series of experi(not shown) indicate this to be a nearly symmetric dipole. An
ments with various eddy size and strength and different
experiment with an anticyclone initially in the lower layer, topographic slope values is being conducted and will be
with the upper layer initially at rest, illustrated that the upper
reported in a separate study.
layer cyclone seen in Figure 7a is the result of vortex
stretching from below, as the lower layer anticyclone dispersesover topography. The experiment in Figure 7 was also
5.
THE EFFECT OF TOPOGRAPHY ON BAROTROPIC
repeated without any coupling to the ice to confirm that the
EDDY-JET INTERACTIONS
dipole formation was not related to the ice cover. A barotroEddy-jet interactions over topography involve combined
pic cyclone (experiment CB) over topographyalso forms an
upper ocean dipole. An anticyclone spun up from below effects of eddy-topography and eddy-jet interactions. Based
appears on the downslope side of the upper ocean cyclone on the eddy-topographyinteractions examined in section 4,
and pairs with it. This dipole pair also propagates to the one might anticipate rapid topographic dispersion in the
"topographic east." The upper and lower layer trajectories lower layer which erodes the lower layer eddy and provides
a strong vortex stretching effect to the upper layer. In the
for this case are also shown in Figure 8.
upper layer, eddy-jet interaction can occur, possibly modiPrevious studies of isolated eddies on a /3 plane [McWilliams and Flierl, 1979; Mied and Lindemann, 1979] have fied by the vortex stretching at the interface. These experishown that initially barotropic eddies in two-mode models ments will show that although the lower layer topographic
decay as barotropic features with eddy motion in each layer dispersionis reduced by the presence of the jet, the vortex
remaining in phase. In those studies, the effect of planetary stretchingeffect on the upper layer is still significantenough
/3 in causing motion and decay is in both layers. Here rapid to modify the upper layer eddy-jet interaction. In addition,
topographic dispersion occurs in the lower layer associated the presence of the topographic vorticity gradient in the
with a topographic /3 effect, but the upper layer equations lower layer can alter the eddy-jet interaction. In the context
have a constant Coriolis field, consistent with the small of the Gulf Stream, Smith and Davis [1989] showed that
planetary /• enhancedcyclone Gulf Stream interaction by
planetary/3 effect at 80øN.
The rapid dispersive decay of the lower layer eddy struc- decreasingthe eddy-jet interaction length scale R. For the
ture in these experiments can be illustrated by considering eastward flowing Gulf Stream however the background
two-layer quasi-geostrophicequations. Following Mied and vorticity gradient enhancesopposite sign eddy-jet vorticity
interaction. For the topographic westward flowing East
Lindemann [1979], in dimensionless form these are written
Greenland Current ice edge jet, topographic /• enhances
like-sign eddy-jet vorticity interaction. Cyclones (anticy-
fo2
2 ]
/•L•J
+•(P2-Pl)
U[Pl,V2pl
fo
2L2 =0
0 V2pl
+•/'H
• 1 (P2-Pl)qOx
Ot
clones)moveupslope(downslope)and have stronger(weaker) interactions with the jet. The topographic eastward
propagating dipoles seen in the upper layer in eddytopographyexperimentsare presentbut lesswell developed
when the jet is included.
SMITH
AND BIRD:
EDDY-JET
INTERACTION
IN A MARGINAL
ICE ZONE
4683
0.75•
0.30•
day0
day2
day4
Fig. 9. Barotropic
cyclone-jet
topography
experiment
CJB, l•'ma
x = 25 cm s-1, days 1, 2, and 4. (a) Ice
concentration
(CI = 0.05);(b)upperlayerrelativevorticity(CI = 0.1 x 10-4 s-l); (c) lowerlayerrelativevorticity
(CI- 0.1 x 10-4 s-l).
5.1. Barotropic Cyclone-Jet Interactions Over
Topography
Upper and lower layer relative vorticity fields show compa-
Figure 9 (experiment CJB) shows the interaction of a
cyclone with the ice-edge jet when a topographic slope is
included. The presence of the jet provides a vorticity barrier
for the free topographic wave dispersionand thus limits the
upslope motion of the lower layer cyclone seen in experiment CB (no jet experiment). The formation of the upper
ocean, "topographic" eastward propagating dipole is thus
slowed (but not eliminated, as is discussedbelow). Because
the dipole formation is slowed, the barotropic to baroclinic
transfer of energy is also reduced. The ratio of available
potential energy (APE) to kinetic energy (KE) on day 4 (see
Table 3) for this experiment is 2 orders of magnitude less
than in the cyclone topography experiment (CB). In comparison to experiment CB, where upper and lower relative
vorticity fields evolve more strongly out of phase, the
eddy-jet topography flow fields evolve more barotropically.
Topography does however alter the eddy-jet interaction.
For cyclones, topography provides an upslope propagation
tendency (Figure 8) which enhancesthe eddy-jet interaction.
A comparison of Figure 8 with Figure 6 indicates that the
cyclone moves further into the ice edge and is advected
rable evolution.
TABLE
3.
APE
Initial
ACJ
CJ
ACB
CB
ACJB
CJB
0
0
0
0
0
0
downstream
morerapidly(6.3 km d-1) than in the flatbottomcase(1.5 km d-l). Thisresultis consistent
with the
findings of Stern and Flierl [1987] who show that downstream motion of like sign vorticies is proportional to 1/R.
Here, topographyhas led to a decreasededdy-jet interaction
length scale R. The trajectories for this case (CJB) and the
flat-bottom case (CJ) are shown in Figure 4.
A secondary circulation associatedwith topographic dispersion in the lower layer is spun up in the upper layer by
vortex stretching at the interface (as in cyclone-topography
case CB) and appears as an anticyclone downslope of the
Energetics
KE1
Final
0.75 x
0.95 x
0.13 x
0.12 x
0.41 x
0.16 x
Initial
104
104
109
109
108
108
0.65 x
0.67 x
0.46 x
0.46 x
0.65 x
0.67 x
KE2
Final
109
109
109
109
109
109
0.48 x
0.50 x
0.19 x
0.13 x
0.44 x
0.54 x
Initial
109
109
109
109
109
109
0.10
0.11
0.23
0.21
0.64
0.66
APE, availablepotential energy;KE1, upper layer kinetic energy;KE2, lower layer kinetic energy.
x
x
x
x
x
x
1012
1012
l0 ll
l0 TM
l0 ll
l0 ll
Final
0.82
0.85
0.11
0.14
0.46
0.56
x
x
x
x
x
x
l0 ll
l0 ll
l0 ll
l0 ll
l0 TM
l0 ll
4684
SMITH AND BIRD: EDDY-JET INTERACTION IN A MARGINAL ICE ZONE
0.75
0.50
day0
day2
day4
Fig. 10. Barotropic
anticyclone-jet
topography
experiment
ACJB,Vma
x = 25cms-1, days1, 2, and4. (a) Ice
concentration
(CI = 0.05);(b)upperlayerrelative
vorticity
(CI = 0.1 x 10-4 s-l); (c)lowerlayerrelative
vorticity
(CI = 0.1 x 10-4 s-l).
cyclone.This anticycloneinteractswith the jet in muchthe
same fashion as in flat-bottom anticyclone-jet experiment
ACJ (Figure 3). It advectscyclonicvorticity away from the
ice edge, creating a second cyclonic ice edge tongue upstream of the original one. The signatureof this in the ice
edge is two cyclonic cusps in the outer ice edge with
approximately25 km spacing.The inner ice edgedoesnot
showthe samedegreeof meandering.It is interestingto note
that the ice edge meanderingis associatedwith the eddies
and not actual meanderingof thejet itself. Meanderingof the
barotropicjet is suppressedby the stabilizingeffect of
topography.
The flat-bottom, cyclone-jet experiment in section3 did
not exhibit any upstreamice edge effects.This experiment
illustratesa possiblemechanismfor creatingan upstreamice
tongue. In this experiment, the second cyclonic ice edge
distortion is associated with the topographic slope. Topographicdispersionin the lower layer cycloneforcesan upper
layer anticyclone;the anticyclonethen interactswith the
cyclonicvorticity of the ice edgejet, advectingice seaward,
upstream of the original cyclone.
5.2. Barotropic Anticyclone-Jet Interactions Over
Topography
In flat-bottom, anticyclone-jet interactions, we have
shown that seaward ice advection can be associated with
The upperoceananticyclonemovesseawardand upstream,
interactingwith cyclonicvorticity which initially appearsto
have comefrom the cyclonicedgeof thejet (as in flat-bottom
caseACJ). In contrastto flat-bottom experimentACJ, where
the cycloneis advectedlaterally away from the jet by the
anticyclone,the cyclone in this experiment is continually
forcedfrom below (as in anticyclone-topographyexperiment
ACB) and remainsupslopeof the anticyclone.Their interaction results in topographic eastward propagation very
comparableto the dipolein experimentACB (Figure7). The
trajectory for this case (experiment ACJB, Figure 10) is
nearlyidenticalto that for ACB (rather than purely seaward
for ACJ) suggestingthat eddy-topographicinteraction is
more importantthan eddy-jet interaction. This is also supportedby an APE/KE ratio (Table 3) which is a factorof 4
greaterin ACB than in the correspondingcyclonecaseCJB,
where the upstreampropagatingdipole is weaker.
While an ice edge distortion initially forms, associated
with the anticyclonebarotropicjet interaction,it is advected
downstreamwith the jet. The growth of this perturbationon
the ice edge is suppressedby topography which has a
stabilizingeffect on the barotropicjet. This effect is further
illustrated in the next section. The ice edge acquires a
cyclonic distortionas the perturbation is advecteddownstream (day 4, Figure 9a).
Figure 4 showsthe trajectoriesfor the flat bottom and the
"dipole" formationin the experimentACJ (Figure3). When slopingbottomeddy-jetcases.It illustratesdifferencesbea topographicslopeis included,this processis inhibitedfor tween anticyclonesand cycloneswith respectto jet-induced
barotropiceddies. For an anticyclone,topographicdisper- motion and differences between flat-bottom and slopingsion results initially in lower layer eddy downslopepropa- bottomexperiments.Motion away from thejet in flat bottom
gation weakening lower layer anticyclone-jet interaction. anticyclonecaseACJ is associatedwith two effects:vortic-
SMITH
AND BIRD'
EDDY-JET
INTERACTION
ity pairing of the eddy with opposite sign vorticity from the
edge of the jet (Figure 3) and azimuthal mode 1 distortion
(Figure 5) acquired by the eddy in the interaction. In cases
with sloping bathymetry, motion is also associated with
topographic wave radiation. Topographic dispersion provides upslope (downslope) motion for the lower layer cyclones (anticyclones). This enhances cyclone-jet interaction
and weakens anticyclone-jet interactions. For cyclones, the
upslope motion results in a strongereddy-jet interaction and
subsequentlyin more rapid downstream advection with the
jet. In the cyclone-jet-topography case CJB, only weak
evidence for the dipole formation associated with eddytopography interaction (CB) is seen. For anticyclones, the
seaward propagation associatedwith eddy-jet interaction is
suppressedand eastward propagation associatedwith eddytopography dipole formation occurs. The azimuthal mode 1
distortion seen in flat-bottom anticyclone-jet experiment
ACJ (Figure 4) is not evident in the correspondingtopography case ACJB.
6.
SENSITIVITY
TO FLOW
STRENGTH
OF TOPOGRAPHIC
AND DEGREE
SLOPE
IN A MARGINAL
ICE ZONE
4685
For a given flow strength, the effect of increased topographic slope is to further reduce ice edge distortions. This is
related to the effect of the topographyon both the jet and the
eddy. Through vorticity constraints, jet perturbations induced by the eddy are inhibited from growing in sloping
bottom experiments. Eddy decay rate is increased for increased topographic slope by rapid topographic dispersion
thus weakening the eddy-jet interaction.
For a given topographicslope, the effect of increasedflow
strength is to enhance ice edge distortions. Stronger eddies
resist dispersive decay [McWilliams and Flierl, 1979] thus
persisting longer, having stronger eddy-jet interactions.
Stronger flows have increased cross-isobath eddy motion
and increased cross-isobathjet meandering.
The seaward extending dipole mechanism associatedwith
flat-bottom anticyclone-jet experiment ACJ is suppressedin
all topographic cases. Dipole mushroom like vortices are
however seen in the ice edge in both Figures 9 and 10 and are
more well defined in high Rossby number experiments.
These dipole shapes can have orientations evolving both
seaward and iceward. A qualitative comparison of some of
these mushroom shapes in Figure 9 with those in Figure 10
indicates
that it can be ditficult
to determine
the sense of
The sensitivity of barotropic eddy-jet interactions to the
degreeof topographicslopein the lower layer is examinedin
this section. In studiesof quasi-geostrophicturbulence over
topography [Rhines, 1976], it has been shown that to first
order, results depend on flow strength and degree of bottom
rotation of the initial eddy from the ice concentration field.
slope.Here the Rossbynumber(Ro = •max/.fL)quantifiesthe
The results of barotropic simulations indicate that topography can significantly modify the interaction of eddies with
the ice edge front. Anticyclone-jet interactions are inhibited
in those experiments and cyclone-jet experiments are augmented. In both cases, the effect of increased topographic
steepnessis to limit cross-ice-edgedistortions and lead to
enhanced eddy decay rate. To investigate the effect of a
baroclinic component on these results, several of the experiments were repeated with vertical shear in the initial con-
strength of the flow and the steepnessof the topography is
indicatedby/J = (l/H2) (Oh2/Oy).For smallRo and large/J,
topography controls the flow and the fluid conservespotential vorticity by following contours of the bottom. For large
values, nonlinearity can dominate and cross-bathymetry
flow is possible. In the context of an isolated eddy over a
sloping bottom, we anticipate these parameters to be an
appropriate quantifier of the results as well. Small Ro, large
/J cases will delineate a more linear regime in which lower
layer flow must conservelinear potential vorticity, dispersing linear topographic Rossby wave energy along slope. For
larger values, the cross-topography self advection seen
above in experiments ACJB and CJB should occur. To test
these hypotheses we have repeated those experiments for
weaker flows and steeper topography, thus varying Ro and
7.
BAROCLINIC
EDDY-JET
INTERACTIONS
OVER TOPOGRAPHY
dition.Upperlayermaximum
velocity(•lmax)is 40 cms-1
andthecorresponding
lowerlayervelocityis 10cms-1 for
both the eddy and the jet in these experiments. These values
for flow strength are consistent with observations in the East
Greenland Current region (J. Johannessen,personal communication, 1989).
In contrast to the barotropic simulations, baroclinic eddyjet interactions(Figures 11 and 12) are largely unmodified by
The steepness (/J) of the continental slope in the East the effect of the slopingbottom. The upper and strongerjet
Greenland
MIZ regionrangesfrom1.5 x 10-5 near76øNto is decoupled from the lower topographically steeredjet and
5.0 x 10-5 near77.5øN.Thusourinitialtopography
results cross-ice-edge perturbations are free to evolve. The lower
apply to the southern location. The effect of steeper topog- layer eddy signatureis rapidly eroded in both anticyclone-jet
raphy is consideredhere. The topographicslopeparameter/J and cyclone-jet experiments, leaving upper layer eddy-jet
is doubled
to 3 x 10-5. Thiscorresponds
to a depthchange interactions to occur as in the flat-bottom experiments. The
from 400 m to 3000 m in 50 km. A flat 400-m-deepcontinental lower layer decay is associated with rapid topographic
shelf extends 30 km from the upper lateral boundary of the Rossbywave radiation along slope. The rate of decay is such
domain. Topographic/3T [-(fo/H2) (Oh2/Oy)]for these two that the lower layer eddy structure is nearly completely
slopevaluesare2 x 10-9 m- 1 s- 1 and4 x 10-9 m- 1 s- l, eroded within several days. Lower layer rotational velocity
decreases
to less than 1 cm s-• in this experiment.In
3 orders of magnitudegreater than planetary/3 at 80øN.
Ice concentration on day 4 for nine barotropic cyclone contrast, the jet lower structure is not affected by this
experiments in which the Rossby number and the topo- process. This evolution to upper layer eddy structure by
graphic slopeparameter/Jare varied in the initial conditionis wave radiation is consistent with the previous studies of
shown in Figure 9, and the corresponding fields for the McWilliams and Flierl [1979] and Mied and Lindemann
anticyclone-jet experiments are shown in Figure 10. The [1979] where the decay of mixed mode eddies on a/3 plane
signature of the eddy-jet interaction in the ice varies sub- were examined. These effects can be seen in relative vorticstantially with both parameters.
ity plots [Figures 13 and 14] where the upper layer vorticity
4686
SMITH AND BIRD: EDDY-JET INTERACTIONIN A MARGINAL ICE ZONE
CJN2
CJN
CJN3
CJB
CJB3
CJ
C J3
0.70
CJB2
I
0.75
0.50
C J2
ß17
.51
.34
Rossbynumber
Fig. 11. Ice concentrationon day 3 for all barotropiccyclone-jetexperiments.
ACJN2
ACJN
ACJN3
ACJB2
ACJB
ACJB3
ACJ
ACJ3 •
0.75
0.45
ACJ2
.17
.34
.51
Rossbynumber
Fig. 12. Ice concentrationon day 3 for all barotropicanticyclone-jetexperiments.
SMITH AND BIRD: EDDY-JET INTERACTIONIN A MARGINAL ICE ZONE
F'
i _
.0.75•
o65
75
0
'
"•
4087
0.70
O.3O
:: .................
0.25
.........
c
day o
day 2
day 4
Fig. 13. Baroclinic
anticyclone-jet
topography
experiment
BCACJB,Vlmax
-- 40 cm S-1, t,'2max
"' 10cm s-1,
days
0,2,
4.s(a)•
Ice
concentration;
(b)upper
layer
relative
vorticity
(CI
-4s-l
day
0,10
0.05
day
2;0.025
xand
10-4
, day
4);(c)lowerlayer
relative
vorticity
(CI
L-0.05
x=
100.5
-4 sx-l,10
day
0;,0.01
x
-4 -4
s-ls-l,
, day
2;
0.005x 10-4 s-1, day4).
0.75
i
i 0 . 30 •
....
day 0
day 2
day 4
Fig. 14. Baroclinic
cyclone-jet
topography
experiment
BCCJB,Vlmax
----40cms-1, V2max
--'-10cms-l, days0, 2,
and4. (a) Iceconcentration;
(b)upperlayerrelative
vorticity
(CI = 0.25x 10-4 s-l , day0; 0.05x 10-4 s-l , day2;
0.05x 10-4 s-l , day4); (c) lowerlayerrelativex;orticity
(CI = 0.05x 10-4 s-l , day0; 0.01x 10-4 s-l , day2;
0.005X 10-4 S-l , day4).
4688
SMITH AND BIRD: EDDY-JET
INTERACTION IN A MARGINAL
extremum associatedwith the eddy is clearly evident on day
4. In contrast, the lower layer vorticity signature of the
original eddy is very weak.
The upper layer responsein these baroclinic experiments
is qualitatively similar to the flat-bottom experiments (section 3). Ice concentrationfields (Figures 13a and 14a) show
a greater degree of spreadingof ice over the axis of the jet
associated with the increased upper layer flow and hence
larger Ekman divergence effect. The growth of downstream
perturbations thus appears to be confined to the outer ice
edge. Anticyclone-jet interactions (Figure 13) produce a
seaward propagating dipole as in flat-bottom experiment
ACJ. The cyclone-jet experiment (Figure 14) produces a
cyclonic cusp in the ice edge but does not provide seaward
ice export.
8.
CONCLUSIONS
The interaction of open ocean eddies with a marginal ice
zone ocean jet has been investigated numerically. The
experiments may help to explain several aspectsof ice edge
distortions seen in remotely sensed images of the East
Greenland Current MIZ. Over a fiat bottom, the interaction
of an initially barotropic anticyclone with the cyclonically
sheared side of the jet can produce dipolar ocean flow fields
and advect ice seawardin a mushroompattern similar to that
seenin syntheticaperture radar images.Observations,however, indicate that the eddiesin the region are predominantly
cyclonic. Cyclonic eddies move toward the jet and are then
advected downstream, creating cyclonic perturbationsin the
ice edge.
Eddy-topography interaction can also cause upper ocean
dipole formations. On short time scales (<1 day), topography induces upslope (downslope) motion in the lower layer
for cyclones (anticyclones). Rapidly evolving lower ocean
topographic wave radiation deforms the interface, providing
strong vortex stretching to upper layer fluid columns. This
strong effect on the upper layer is a consequenceof the
thinnessof the upper layer (50 m) relative to the lower layer
(4000 m). Upper layer dipoles form from both anticyclones
and cyclones and propagateupstream, alongslopewith shallow water to the left (topographic east), in both cases. The
lower layer structures evolve rapidly out of phase with the
upper layer dipole, with the lower layer anticyclone
(cyclone) moving downslope (upslope). The results suggest
that barotropic eddies advected into the East Greenland
continental slope region are unlikely to have deep structure
after even brief interactions with the sloping bottom. The
effect of topography on barotropic eddy-jet interactionsis to
enhancecyclone-jet interactions and weaken anticyclone-jet
interactions. Topography also tends to stabilize the jet,
limiting downstreamgrowth of eddy-inducedperturbations
on the jet. Cross-isobathmotions in the meanderingice edge
jet are only seen in high Rossby number experiments.
Baroclinic eddy-jet interactions over topography, however, more closely resemble flat-bottom experiments, suggesting that topography may not be important in baroclinic
eddy-jet interactions. Rapid topographic wave radiation
erodesthe lower layer eddy structure,leavingupper layer
eddy-jet interactions which evolve independentof the effect
of topography. In contrast, the lower layer of the jet is not
weakened by this process.The results of flat-bottom experiments then suggestthat anticyclone-jet interactions will
ICE ZONE
produce seaward-extending dipolar ice flow fields within
time scalesof several days for upper ocean velocities in the
rangeof 10to 40 cm s-1 (Ro = 0.17-0.54).The extentof
seawardice export increaseswith flow strength.Cyclone-jet
interactions do not provide seaward dipolar flows in the ice
but lead to downstreamcyclonic-shapedice tongues.These
experiments also illustrate that it can be difficult to determine the sense of eddy rotation or number of eddies from
spatial patterns in the ice concentrationfield.
NOTATION
A
ice concentration.
A0 eddy maximum amplitude.
A h Laplacian lateral friction coefficient,equal to 10
m 2 s-1 '
A a Laplacian diffusioncoefficientfor ice
concentration, equal to 30.
A m Laplacian diffusioncoefficientfor ice mass,
equal to 30.
/• variation of Coriolis parameter (80øN) with
latitude,equalto 3.8 x 10-12 m-1 s-•.
Ciw ice-water interfacial stresscoefficient,equal to
7.5 x 10-3.
d
D
variable depth topography.
ice thickness,equal to rn/piA.
•x
At
grid spatial resolution, equal to 1 km.
finite difference time step, equal to 600 s.
•il
Kronecker delta (equal to 0 when i = 2).
V2 Laplacianoperator,equalto (02/0x2 + 02/0y2).
f0
Coriolis parameterfor mean latitude (80øN),
equalto 1.43x 10-4 s-1.
•7 acceleration
dueto gravity,equalto 9.8 m s-2.
•7' reducedgravity for upper, lower density Pi,
equalto 0.02m s-2 (= [(P2-- Pl)g/Pl])'
•,
Hi
H2
hi
/max
nondimensionaleddy size, equal to L/Ra.
upper layer mean thickness, equal to 50 m.
lower layer mean thickness,equal to 4000 m.
instantaneouslayer thickness.
number of grid points in the zonal direction,
equal to 111.
Jmax number of grid points in the meridional
direction, equal to 79.
L e e-foldingscalefor the eddy, equal to 5 km.
Lj e-foldingscalefor thejet, equalto 5 km.
L(0)
I
rn
P•
initial distance from the eddy center to the jet
edge.
eddy interaction length scale, equal to 9 km.
ice mass.
pressurein the upper layer, equal to 9(hl + h2
+ d).
P2 pressurein the lower layer, equal to P1 - 9'hl.
Q nondimensionaleddy strength,equal to •'max/
qi upper, lower layer potentialvorticity, equal to (f
+ •i)/hi.
R d first internal Rossbyradiusof deformation,equal
to (l/f) [(•I'H•H2)/H•H2]
m.
Roi
r
Rossby number for upper (i = 1) and lower
(i = 2) layers, equal to V/max/fL.
eddy radius, equal to 5 km.
p• density
of ice,equalto 910kgm-3.
riw ice-waterinterfacial
stress,equalto plCiw
SMITH
AND BIRD: EDDY-JET
INTERACTION
u•, v• ice velocities in the x, y directions.
uw, vw ocean upper layer velocities in the x, y
directions.
Ui, Vi
ICE ZONE
4689
layer quasigeostrophicmodel, in Mesoscale/Synoptic Coherent
Structures in Geophysical Turbulence, edited by J. C. J. Nihoul
and B. M. Jamart, pp. 277-291, Elsevier, New York, 1989.
Johannessen, J. A., O. M. Johannessen, E. Svendsen, R. Shuch-
man, T. Manley, W. J. Campbell, E.G. Josberger, S. Sandven,
ocean transport in the x, y directions in layer i 1,2.
Vemax ocean eddy maximum tanõentialvelocity.
Vjmaxoceanjet maximumvelocity.
x, y
•i
IN A MARGINAL
Cartesian coordinates.
upper, lower layer relative vorticity - V x vi.
J. C. Gascard, T. Olaussen, K. Davidson, and J. Van Leer, Mesoscale eddies in the Fram Strait marginal ice zone during the 1983
and 1984 Marginal Ice Zone Experiments, J. Geophys. Res., 92,
6754-6772, 1987.
Johannessen,O. M., J. A. Johannessen,J. Morison, B. A. Farrelly,
and E. A. Svendsen,Oceanographicconditionsin the marginal ice
zone north of Svalbard in early fall 1979 with an emphasis on
mesoscaleprocesses,J. Geophys. Res., 88, 2755-2769, 1983.
Louis, J.P., and P. C. Smith, The development of the barotropic
radiation field of an eddy over a slope, J. Phys. Oceanogr., 12,
Acknowledgments. Support for this work has come from the
56--73, 1982.
Office of Naval Research, Arctic Programs.Computing resources
from the W.R. Church Computing Center are also gratefully ac- McWilliams, J. C., and G. R. Flierl, On the evolution of isolated
nonlinear vortices, J. Phys. Oceanogr., 9, 1155-1182, 1979.
knowledged. Helpful comments were received from J.C. Gascard,
Mied, R. P., and G. J. Lindemann, The propagationand evolution of
A. Foldvik, and R.P. Mied during the course of this research.
cyclonic Gulf Stream rings, J. Phys. Oceanogr., 9, 1183-1206,
1979.
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A. A. Bird and D.C. Smith, IV, Department of Oceanography,
Naval Postgraduate School, Monterey, CA 93943.
(Received April 9, 1990;
revised July 11, 1990;
accepted June 7, 1990.)