PUBLICATIONS
Geophysical Research Letters
RESEARCH LETTER
10.1002/2014GL062755
Key Points:
• Tidal processes in the deep ocean
reflect subkilometer topographic scales
• Observations at these scales reveal
tidally pulsed overflows and lee waves
• Dissipation levels suggest a widely
distributed mixing contribution to
the MOC
Correspondence to:
A. C. Dale,
[email protected]
Citation:
Dale, A. C., and M. E. Inall (2015), Tidal
mixing processes amid small-scale,
deep-ocean topography, Geophys. Res.
Lett., 42, 484–491, doi:10.1002/
2014GL062755.
Received 5 DEC 2014
Accepted 24 DEC 2014
Accepted article online 5 JAN 2015
Published online 30 JAN 2015
The copyright line for this article was
changed on 20 JAN 2016 after original
online publication.
This is an open access article under the
terms of the Creative Commons
Attribution License, which permits use,
distribution and reproduction in any
medium, provided the original work is
properly cited.
DALE AND INALL
Tidal mixing processes amid small-scale,
deep-ocean topography
Andrew C. Dale1 and Mark E. Inall1
1
Scottish Association for Marine Science, Scottish Marine Institute, Oban, UK
Abstract The nature of tide-topography interaction reflects the topographic scales experienced by water
parcels during their tidal excursions. In the deep ocean these scales are typically subkilometer, yet direct
observations of tidal processes on such scales are lacking. At one site, a saddle amid steep and complex
Mid-Atlantic Ridge topography, observations reveal tidally pulsed, bottom-trapped fronts, overflows, and lee
waves in response to a tide combined with a mean flow of similar amplitude. The tidal pulsing of the
fronts and overflows was only evident locally, and their phase became unpredictable over scales of hundreds
of meters. Enhanced turbulence in a 100–200 m thick bottom boundary layer had an estimated dissipation
rate of 2.6 × 102 W m2, exceeding the large-scale average of tidal dissipation in mid-ocean ridge
environments but by less than an order of magnitude. This site was not a dissipation “hotspot,” and the
processes observed could provide widely distributed mixing to the meridional overturning circulation.
1. Introduction
Mixing in the deep ocean draws buoyant water downward. This provides potential energy to the water column
which, in part, sustains the meridional overturning circulation (MOC). Since tides are a key agent of deep mixing
[Munk and Wunsch, 1998; Wunsch and Ferrari, 2004], the magnitude and distribution of tidal mixing profoundly
influence the structure and timescales of the MOC [Hasumi and Suginohara, 1999; Jayne, 2009], and consequently
influence the ability of the deep ocean to buffer the climate system [Meehl et al., 2011].
Tidal currents encountering sloping topography lose energy, both to dissipative processes that drive local
mixing, and to radiated internal waves. In view of the sensitivity of the MOC to the distribution of mixing, it is
important to understand the balance between local dissipation and radiation and the location and efficiency
of their ultimate delivery of energy to mixing. These aspects are highly process dependent and poorly
understood. A key difficulty is that tidal processes reflect topography on the scales experienced by water
parcels during their tidal excursions, and, in the deep ocean, these scales are typically less than a kilometer.
Abyssal topography is rich in these scales, with mid-ocean ridges in particular providing vast tracts of
complex, multiscale topography. While observational confirmation of the role of tidal mixing in mid-ocean
environments comes from a spring-neap modulation of dissipation [Ledwell et al., 2000], direct, process level
observations in such environments are restricted to topographic features that are considerably larger than the
tidal excursion [Alford et al., 2014; Klymak et al., 2008; Van Haren et al., 2005] or cases where the mean flow
dominates the tide [Alford et al., 2013; St Laurent and Thurnherr, 2007; Thurnherr et al., 2005]. An additional
difficulty is that ocean bathymetry is not consistently known at subkilometer scales. Satellite altimetry, while
providing global bathymetric coverage, resolves only scales greater than 25 km [Smith and Sandwell, 1997] and
higher resolution ship-based surveys are only patchily available. Merged best available topographies, therefore,
significantly underestimate both bottom slope and excursion-scale complexity [Becker and Sandwell, 2008;
Zilberman et al., 2009]. The likely significance of unresolved scales in tidal energy conversion [Melet et al., 2013]
poses a challenge for global models of tidal dissipation [Nycander, 2005; Simmons et al., 2004], which also
face computational limits on their resolution. While such models appear to be converging on the known
(astronomically derived) energy loss from the tide, they potentially remain deficient at the process level and
vary in the location of their delivery of energy to mixing [Green and Nycander, 2013].
For an isolated topographic feature, the processes governing tide-topography interaction depend on the
slope of the flanks of the feature, the vertical density gradient, geographic latitude (via the Coriolis parameter f),
and the horizontal scale of the topography relative to the tidal excursion [Garrett and Kunze, 2007]. The
complexity of the situation is compounded if there is nontidal flow or three-dimensionality of background flows
©2015. The Authors.
484
Geophysical Research Letters
10.1002/2014GL062755
Figure 1. Bathymetry of the study site. (a) Overview. (b) Detailed local bathymetry from multibeam echosounder.
(c) Bottom slope relative to the angle of internal wave propagation at the M2 frequency. White = less steep (subcritical),
orange = steeper (supercritical), red = more than 5 times steeper. Bathymetry is contoured (50 m interval). A = thermistor
chain. B = lander. C = lander site C1, with yoyo site C2 located 90 m to the NW. The M2 (lunar semidiurnal) tidal ellipse
(TPXO7.2 model) is shown at site A, scaled to show 5 times the tidal excursion.
or topography. When topographic scales are large compared to the tidal excursion, individual water parcels
experience the topography as a slope. This case is relatively well understood, although turbulent and complex
behavior may occur. When the bottom slope is near-critical (i.e., close to the internal tide propagation angle),
near-bed flow becomes highly sheared and upslope-propagating bores, resembling gravity currents, may be
generated [Aucan et al., 2006; Gayen and Sarkar, 2011; Gemmrich and Van Haren, 2001]. Similar bores have also
been observed at noncritical slope angles [Van Haren, 2005; Van Haren et al., 2005] and may result from internal
waves impinging on a slope [Hosegood and van Haren, 2004].
When topography contains scales that are comparable to or shorter than the tidal excursion, water parcels
experience changes of gradient, permitting processes such as lee waves, hydraulically controlled flows and
overflows. Although changes of gradient on the flanks of larger features are included in this case [Klymak
et al., 2008; Legg and Klymak, 2008], the focus here is on features that are themselves of a size that is comparable
to the excursion scale. While several studies have considered tidal interactions with excursion-scale topography
from a modeling perspective [Legg and Huijts, 2006; Nikurashin and Legg, 2011], detailed observations have
been lacking.
Our studies were conducted on the eastern flank of the Mid-Atlantic Ridge at 49°N (Figure 1) at a site where, in
2007, a single-day lander record had shown the passage of two cold fronts separated by a tidal period. The
detailed setting, revealed by a multibeam bathymetric survey [Niedzielski et al., 2013], was a saddle, the low
point (depth 2535 m) of a spur emanating from a more extensive elevated area to the west and NW (summit
depth 1555 m, 2 km west of the boundary of Figure 1b), set within an extensive tract of complex, mid-ocean
ridge topography. The flanks of the spur were asymmetric, with a short slope (70 m vertical, 250 m horizontal)
to a small basin to the SW, and a longer slope (270 m vertical, 1000 m horizontal) to a more extensive deep
DALE AND INALL
©2015. The Authors.
485
Geophysical Research Letters
10.1002/2014GL062755
Figure 2. Time series during yoyo profiling on 9 August 2009 at C2. (a) Horizontal velocity vectors (y axis is oriented to
north), averaged over the bottom 500 m, from lowered ADCP. (b) Vertical temperature structure. Black dots indicate
Thorpe overturning scales exceeding 25 m. The vertical structure of the tidally averaged TKE dissipation rate (ε) is shown to
the right. (c) Temperature at 4 m above the bed from the bottom lander at nearby site C1. (d) Velocity at 8 m above the bed
from the bottom lander at site C1. Times are relative to low water at A.
area to the NE. Time series from seafloor lander deployments, a moored thermistor chain, and “yoyo” profiles
provide insights into the tidal processes of this site and their spatial variability over scales of hundreds of
meters. These observations allow an estimate of local mixing energetics which will be related to the global
tidal mixing budget.
2. Methods
2.1. Moored Instruments: Thermistor Chain and Lander
Data collection was during August 2009 from RRS James Cook at three sites within 500 m of the saddle
(Figure 1c). At site A, 230 m to the NE of the saddle and 25 m downslope from it, a vertical array of
thermistors (RBR XR420TiD + T16) was moored for 18 semidiurnal tidal cycles. The array consisted of 16
equally spaced thermistors positioned 9–109 m above the seabed and sampling with a 5 s interval. While
the thermistor chain was at A, a concurrent set of observations was made on the opposite flank of the
spur. A bottom lander was deployed, initially for 5–6 tidal cycles at site C1, 370 m to the SSW of the saddle
and 80 m deeper than it. The lander was then redeployed for a further 14 tidal cycles at B, 15 m below
the saddle and 370 m to the WNW of it on the SW flank of a slightly higher part of the spur. Lander
instrumentation consisted of a Seabird 37 CTD (conductivity-temperature-depth) logger sampling at 3.4
m above the bed, and an upward looking 300 kHz Acoustic Doppler Current Profiler (ADCP) recording
velocity profiles in 2 m vertical bins from 8 m to approximately 50 m above the bed. Both instruments
had a 10 s sample interval. In view of the unknown lateral displacement due to horizontal currents as
these moored instruments descended through the water column, their final positions on the seabed were
triangulated acoustically to within an estimated 50 m. Most (10 of the 16) thermistors showed time-varying,
monotonic calibration drift, which has been corrected by comparison with nearby stable thermistors using a
spline temporal correction to match a vertically interpolated mean temperature.
2.2. Yoyo Temperature, Salinity, and Velocity Profiles
Concurrent with the thermistor array at A and the lander at C1, repeated vertical yoyo profiles of temperature,
salinity, and velocity were collected from the ship stationed at site C2, 90 m to the NE of C1, for 24 h using a
24 Hz Seabird 911+ CTD (conductivity-temperature-depth) system. C1 and C2 were essentially collocated
within the known accuracy of their locations, but not necessarily with respect to sub-100 m flow structures.
DALE AND INALL
©2015. The Authors.
486
Geophysical Research Letters
10.1002/2014GL062755
A total of 77 profiles were obtained,
spanning from a depth of 2150 m to
within 10 m of the bed. A pair of
upward and downward looking
300 kHz Acoustic Doppler Current
Profilers (ADCPs) was mounted on the
CTD frame. The vessel held position
using Dynamic Positioning during this
period. Depth-averaged velocities
have been calculated under the
assumption that when bottom
tracking was not available there
was no horizontal motion of the
instrument package. This leads to
an error with an estimated RMS
amplitude of 0.8 cm s1, an order
of magnitude smaller than the
observed variability. The sea state
was slight.
2.3. Dissipation Estimation
From Overturns
Figure 3. Thermistor chain time series of temperature in the bottom 109 m of
the water column at site A, 7–16 August 2009. Time is in hours relative to local
low water for the M2 tidal constituent (TPXO7.2 model). Inset numbers give
the sequential tidal cycle within the data set. Only those tidal cycles where
sharp bottom fronts were observed are shown, with cycle 15 included as
representative of the cycles on which a sharp front was not observed.
Using calibrated 24 Hz temperature
profiles from the yoyo, a conservative
method was used to calculate
the vertical extent of overturns of
the water column [Thorpe, 1977].
Setting the noise level of the
Seabird thermistor to 0.002°C and
defining the Thorpe scale (LT ) as the
root-mean-square of the vertical
displacements of the monotonically
reordered temperature profiles, the
Ozmidov scale (LO) was determined
[Ferron et al., 1998] according to
Lo = 0.95LT and the turbulent kinetic
energy dissipation rate ε = hN3iLo2.
The angle brackets here denote an
average over turbulent (overturning)
patches and N is the buoyancy
frequency of the reordered
water column.
3. Observed Tidal Dynamics
In the central North Atlantic, the tide is dominated by the lunar semidiurnal (M2) constituent. The predicted
depth-mean M2 flow at the study site (TPXO7.2 tidal model [Egbert and Erofeeva, 2002]) is largely east-west
with current ellipses having a major axis of 3.6 cm s1 and a tidal excursion of 506 m (Figure 1). The TPXO
model does not resolve topography smaller than its quarter-degree resolution, however, and local steering of
tidal flows is expected. The observed velocity, averaged over the bottom 500 m of the water column during
the yoyo at C2, showed a tidally pulsed flow peaking at 7–8 cm s1 to the SW (Figure 2a). This flow largely
stalled on each tidal cycle but did not reverse to the NE, a pattern consistent with tidal currents comparable in
amplitude to the TPXO predictions superimposed on a background flow to the SW of similar magnitude.
DALE AND INALL
©2015. The Authors.
487
Geophysical Research Letters
10.1002/2014GL062755
Two distinct phenomena were
observed in response to this tidal
forcing: the pulsing of dense
water across the saddle during
the strongest flow to the SW,
and the development and
release of internal lee waves.
Bottom-trapped pulses of
relatively cool (dense) water
(Figure 3) at the thermistor array
site, A, traveled to the SW across
the saddle (the direction is
inferred from the near-bed
current associated with these
features in the 2007 observations).
Of these cool pulses, 10 of 18
were led by a sharp front at which
near-bed temperatures dropped
by at least 0.05°C in 10 min,
nearly half of the total observed
temperature range at this site and
Figure 4. Near-bed temperature time series at sites A, B, and C1. Time wraps in equivalent to a density increase of
tidal phase and is expressed as hours relative to the local low water for the M2
approximately 0.01 kg m3
tidal constituent (TPXO7.2 model). Sharp temperature drops (>0.05°C in <10 min) (temperature and density were
are shown in red. Instrument heights above the bed are 9 m (A), 4 m (B), and 4
related monotonically). On no
m (C1). Data periods are 7–16 August (A), 7–10 August (B), and 11–18 August
occasion was there a similarly
(C1) 2009.
sharp rise in temperature. The
cool pulses were clearly tidal, with
sharp fronts clustering in phase
between 5.1 and 7.3 h after local low water (Figure 4) as the N-S flow at C2 stalled prior to building to the
SW (Figure 2a). The fronts were traveling upslope, approaching the crest of the spur. No such fronts were
observed on the opposite phase of the tide. The vertical and temporal structure of the fronts varied
considerably (Figure 3), with their thickness ranging from O(10 m) to more than 100 m. Typical
characteristics were a sloping leading edge and an elevated head with overturning on its trailing side
followed by a depression and gradual thickening of the cool layer. In these respects they strongly
resembled the upslope-propagating bores described by Van Haren [2005] and modeled by Gayen and
Sarkar [2011]. There was no clear spring-neap cycle to the occurrence, timing, or structure of these fronts
(spring tides occurred early in the record and neap tides toward the end). The role of slope criticality in their
formation is unknown; slopes in the region are largely supercritical (steeper than the internal tide
propagation angle; Figure 1c) with transitions through criticality on the flanks of summits and troughs,
including near A. The tidal phasing at A suggests generation downslope of this site on the NE flank of the
spur. The tidal asymmetry of this process (cool pulses pass over the saddle to the SW, but not to the NE on
the opposite phase of the tide) is likely imposed by the background flow; however, topographic asymmetry
may also play a role.
Following the passage of each front, cool water persisted at A for several hours during the strongest flow to
the SW at C2. It is assumed that each cool pulse continued for the 230 m to the summit and spilled down
the slope to the SW as a gravity current with an estimated frontal speed [Shin et al., 2004] of 0.07 m s1
pffiffiffiffiffiffiffi
(c ¼ g′H where g ′ = gΔρ/ρ is the reduced gravity, cool layer thickness H = 50 m, frontal density step
Δρ = 0.01 kg m 3, ambient density ρ = 1040 kg m 3, and g is the acceleration due to gravity). This speed is
comparable to the peak depth-mean current. An estimate of the resultant energy flux E across the spur can
be obtained from the sum of kinetic and potential energy fluxes based on water of depth H with density
anomaly Δρ traveling at speed c. The kinetic energy density of this flow is ρc2/2 and its potential energy density
DALE AND INALL
©2015. The Authors.
488
Geophysical Research Letters
10.1002/2014GL062755
is ρg ′ 2/2N2 under the assumption that it has been raised from ambient stratification with a buoyancy frequency
of N = 10 3 s 1 (based on nearby CTD profiles). The energy flux across the spur is then
ρ
g′
E ¼ c3 H þ 2 ¼ 25 W m1 :
2
N
The basin receiving this energy extends for around 1 km normal to the spur, comparable to the internal
Rossby radius over which rotational adjustment is expected (Ro = c/f = 629 m in this case) [Shapiro and
Zatsepin, 1997]. Assuming that the system is two-dimensional, that the energy flux is constant, and that
all energy is dissipated within 1 km of the ridge yields a mean vertically integrated dissipation rate of
2.5 × 102 W m2. This is clearly complicated by three dimensionality and reduced in a tidally averaged sense
by the intermittency of the process.
Lander records from the SW flank of the spur (sites B and C1, Figure 1c) also show the passage of cool fronts,
with a frequency of occurrence that is comparable to those at site A. These fronts, however, are less consistent
in their tidal phase (Figure 4). Breakdown of phase-locking with distance from the source is expected due
to variation of structure and propagation speed between tidal cycles as well as three-dimensional complexity
of the system.
The yoyo profiles at C2 provide a glimpse of tidal dynamics higher in the water column (Figure 2b).
Temperature structure here reveals vertical displacements with quarter-diurnal periodicity dominating
between 500 m and 200 m above the bed (Figure 2b, Q1 to Q4). A water column advected semidiurnally back
and forth across a summit is displaced vertically with a quarter-diurnal frequency; therefore, tidal interaction
with excursion-scale topography has a natural tendency to radiate quarter-diurnal internal waves. Flow
reversal is required, however, so the observed quarter-diurnal dominance suggests that flow was not
everywhere unidirectional as suggested by the depth-averaged currents at C2 (Figure 2a). The vertical extent
of oscillations of the water column was around 100 m, comparable to the height of the flanks of the spur.
Within 200 m of the bed, stratification was weak and density structure was frequently inverted. Peak flow to
the SW led to a drawing down of lighter water toward the bed in the lee of the spur. This displacement
was most pronounced around 1 h after low water (Figure 2b). As flow weakened, this light water rapidly
rebounded and a cold bottom front passed the nearby lander at around 2 h after low water on each tidal
cycle (Figure 2c). Near-bed currents following the passage of the front were to the north (Figure 2d). These
features are consistent with the formation of a steep lee wave and hydraulic jump during peak downslope
flow, followed by upslope propagation of these structures as flow weakens. This is a different mechanism
from that responsible for the fronts observed at A.
Lee wave formation occurs when flow over a topographic feature stalls the propagation of internal perturbations.
In a strengthening ambient flow U, high vertical modes are the first to be stalled and there is a tendency
for upstream flow to be blocked beneath the level of the topography [Klymak et al., 2010]. As the flow speed
increases, the lowest mode stalls at a topographic Froude number Fr = U/Nh ≈ 1 (where h is the height of the
topography). Lee waves then steepen to the point of overturning [Baines, 1995] and form a downstream
hydraulic jump. In the present case, Fr = 1 is realized at near-peak flow U = 0.07 m s 1 for an ambient buoyancy
frequency N = 10 3 s 1 and h = 70 m, representing the SW flank of the spur. In an oscillating tidal flow, an
additional requirement for the formation of a hydraulic jump is that the slope is steep (supercritical) [Legg and
Klymak, 2008], as in the present case. Observations to date of tidal lee waves and hydraulic jumps in the deep
ocean have been associated with significantly taller topography than is considered here [Alford et al., 2014;
Klymak et al., 2008; Pinkel et al., 2012], combined with stronger flows and smaller Froude number. The present
observations provide confirmation of lee wave formation due to small-scale topography, a situation which has
previously been restricted to numerical studies [Legg and Huijts, 2006].
Overturns larger than 25 m, corresponding here to a dissipation rate ε > 2.0 × 107 W kg1, reveal an
approximately semidiurnal periodicity to mixing centered around the 3.22°C isotherm (~250 m above the
bed, W1 and W2 in Figure 2b), and independently within the 100 m to 200 m thick bottom layer (B1 to B3,
Figure 2b). B1 and B2 persist for several hours and are associated with peak flows and lee wave formation and
release. B3 occurs earlier in tidal phase, is more strongly bottom trapped, and is associated with a patch of
cool water in the bottom tens of meters with a sharp leading front (Figure 2c). Passage of the front is followed
by near-bed flow to the south (Figure 2d). We cautiously identify this as a downslope-propagating gravity
DALE AND INALL
©2015. The Authors.
489
Geophysical Research Letters
10.1002/2014GL062755
current. W1 and W2 occur higher in the water column, with near-opposite tidal phasing to B1 and B2.
Although three-dimensional complexity of the site is such that the source of these features is uncertain,
they resemble patches of vertical straining and mixing that can result from lee wave formation and release
[Klymak et al., 2008].
Vertical mixing and dissipation associated with the observed locally intensified tidal phenomena were
large. Peak TKE (Turbulent Kinetic Energy) dissipation rates, derived from the largest individual overturns,
were ε = 2.1 × 106 W kg1 within the bottom features B1 to B3 and ε = 2.5 × 106 W kg1 within the water
column features W1 and W2. Averaging over two tidal cycles gives a depth-integrated dissipation rate
of 2.6 × 102 W m2.
While the spatial complexities of sites such as this cannot be unraveled from a small number of sampling
locations, the observations provide insight into the phenomena to be expected in such environments
and their associated energetics. In particular, they emphasize the fact that the lateral scales of variability
reflect the subkilometer scales of the tidal excursion, the internal Rossby radius, and similar topographic
scales.
4. Upscaling
Astronomical observations place the dissipation of tidal energy in the ocean at around 3.5 TW [Munk and
Wunsch, 1998]. Satellite altimeter estimates suggest that, of this, around 1 TW is lost by the tides in the deep
ocean as a result of topographic interactions [Egbert and Ray, 2001], contributing nearly half of the 2.1 TW
required to support the MOC [Munk and Wunsch, 1998; Wunsch and Ferrari, 2004].
Of the 1 TW of tidal energy that is lost in the deep ocean, estimates of the locally dissipated component
vary widely [Munk and Wunsch, 1998; St Laurent and Garrett, 2002]. Assuming that 0.7 TW of the 1 TW is
dissipated locally [Munk and Wunsch, 1998] and that this dissipation is evenly distributed over the 23% of
the Earth’s surface that is covered by mid-ocean ridges [Heezen, 1969] yields a mean spatial density of local
tidal dissipation of 6 × 103 W m2. This is around a quarter of the dissipation that we have estimated
from the bottom 500 m of the yoyo and half the dissipation estimated from gravity current energetics
downslope of the saddle. Dissipation at this site is therefore elevated above the expected large-scale
spatial mean, but not by orders of magnitude. While dissipation “hot spots” have been identified in several
deep-ocean locations [Alford et al., 2013; MacKinnon et al., 2008; Pinkel et al., 2012], such sites are, by
definition, atypical. In contrast, less dissipative phenomena, such as those described here, could be widely
present in complex bathymetric environments, setting the fluid dynamical character of more typical
mid-ocean ridge sites.
5. Conclusions
Currents tend to be weak in the deep ocean. In this weakly stratified environment, however, even
relatively low-energy flows may be highly turbulent. Tidal processes are expected to reflect the
subkilometer scales of the tidal excursion, of the internal Rossby radius, and similar scales in the seafloor
topography. Attempts to observe tidal physics on these scales, however, are confounded by inherent
complexity, variability and three dimensionality as well as by the practical problems of observing the deep
ocean. These include the difficulty of accurately positioning instrumentation to sub-100 m accuracy
from a vessel several kilometers above. Nevertheless, we have described tidally pulsed fronts, overflows,
and lee waves from an essentially randomly selected site. The tidal phase locking of these phenomena,
while clear near to their generation site (a saddle), was rapidly lost over just a few hundred meters. These
were very local phenomena which did not always clearly reveal their tidal origin. The ubiquity of such
topography suggests that such phenomena may be widely present in deep-ocean environments but
also site specific in their character, reflecting local topography and the relative local importance of tidal
and background low-frequency flow. While global general circulation models are far from resolving
topography and physics on the tidal excursion scale, it is suggested that parameterizations of tidal
dissipation and mixing would better reflect process-level physics if they reflected regional differences in
the small-scale character and complexity of topography.
DALE AND INALL
©2015. The Authors.
490
Geophysical Research Letters
Acknowledgments
All data sets described are lodged with
the British Oceanographic Data Centre.
This work was funded by the UK Natural
Environment Research Council as part
of the ECOMAR consortium, grant
number NE/C512961/1. We are grateful
to the consortium leader, Monty Priede,
for his support of this aspect of the
study, and to the lander team at
Oceanlab, University of Aberdeen,
for enabling us to make physical
observations from their lander platform.
Colin Griffiths and John Beaton
provided invaluable assistance with
observational aspects of the study,
and Tomasz Niedzielski processed the
multibeam bathymetric data set.
Preparation of this manuscript was
completed with funding from the
European Union Seventh Framework
Programme (FP7/2007-2013) under the
MIDAS project, grant agreement 603418.
The Editor thanks two anonymous
reviewers for their assistance in evaluating
this paper.
DALE AND INALL
10.1002/2014GL062755
References
Alford, M., J. B. Girton, G. Voet, G. S. Carter, J. B. Mickett, and J. M. Klymak (2013), Turbulent mixing and hydraulic control of abyssal water in
the Samoan Passage, Geophys. Res. Lett., 40, 4668–4674, doi:10.1002/grl.50684.
Alford, M., J. M. Klymak, and G. S. Carter (2014), Breaking internal lee waves at Kaena Ridge, Hawaii, Geophys. Res. Lett., 41, 906–912,
doi:10.1002/2013GL059070.
Aucan, J., M. A. Merrifield, D. S. Luther, and P. Flament (2006), Tidal mixing events on the deep flanks of Kaena Ridge, Hawaii, J. Phys. Oceanogr.,
36(6), 1202–1219.
Baines, P. G. (1995), Topographic Effects in Stratified Flows, Cambridge Univ. Press, Cambridge, U. K.
Becker, J. J., and D. T. Sandwell (2008), Global estimates of seafloor slope from single-beam ship soundings, J. Geophys. Res., 113, C05028,
doi:10.1029/2006JC003879.
Egbert, G. D., and S. Y. Erofeeva (2002), Efficient inverse Modeling of barotropic ocean tides, J. Atmos. Oceanic Technol., 19(2), 183–204,
doi:10.1175/1520-0426(2002)019<0183:eimobo>2.0.co;2.
Egbert, G. D., and R. D. Ray (2001), Estimates of M2 tidal energy dissipation from TOPEX/Poseidon altimeter data, J. Geophys. Res., 106(C10),
22,475–22,502, doi:10.1029/2000JC000699.
Ferron, B., H. Mercier, K. Speer, A. Gargett, and K. Polzin (1998), Mixing in the Romanche fracture zone, J. Phys. Oceanogr., 28(10), 1929–1945.
Garrett, C., and E. Kunze (2007), Internal tide generation in the deep ocean, Annu. Rev. Fluid Mech., 39, 57–87.
Gayen, B., and S. Sarkar (2011), Direct and large-eddy simulations of internal tide generation at a near-critical slope, J. Fluid Mech., 681, 48–79.
Gemmrich, J. R., and H. Van Haren (2001), Thermal fronts generated by internal waves propagating obliquely along the continental slope,
J. Phys. Oceanogr., 31(3), 649–655.
Green, J., and J. Nycander (2013), A comparison of tidal conversion parameterizations for tidal models, J. Phys. Oceanogr., 43, 104–119.
Hasumi, H., and N. Suginohara (1999), Effects of locally enhanced vertical diffusivity over rough bathymetry on the world ocean circulation,
J. Geophys. Res., 104(C10), 23,367–23,374, doi:10.1029/1999JC900191.
Heezen, B. C. (1969), The world rift system: An introduction to the symposium, Tectonophysics, 8(4–6), 269–279, doi:10.1016/0040-1951(69)
90037-7.
Hosegood, P., and H. van Haren (2004), Near-bed solibores over the continental slope in the Faeroe-Shetland Channel, Deep Sea Res., Part II,
51(25–26), 2943–2971, doi:10.1016/j.dsr2.2004.09.016.
Jayne, S. R. (2009), The impact of abyssal mixing parameterizations in an ocean general circulation model, J. Phys. Oceanogr., 39(7), 1756–1775.
Klymak, J. M., R. Pinkel, and L. Rainville (2008), Direct breaking of the internal tide near topography: Kaena Ridge, Hawaii, J. Phys. Oceanogr.,
38(2), 380–399, doi:10.1175/2007jpo3728.1.
Klymak, J. M., S. M. Legg, and R. Pinkel (2010), High-mode stationary waves in stratified flow over large obstacles, J. Fluid Mech., 644, 321–336,
doi:10.1017/s0022112009992503.
Ledwell, J., E. Montgomery, K. Polzin, L. S. Laurent, R. Schmitt, and J. Toole (2000), Evidence for enhanced mixing over rough topography in
the abyssal ocean, Nature, 403(6766), 179–182.
Legg, S., and K. M. Huijts (2006), Preliminary simulations of internal waves and mixing generated by finite amplitude tidal flow over isolated
topography, Deep Sea Res., Part II, 53(1), 140–156.
Legg, S., and J. Klymak (2008), Internal hydraulic jumps and overturning generated by tidal flow over a tall steep ridge, J. Phys. Oceanogr.,
38(9), 1949–1964, doi:10.1175/2008jpo3777.1.
MacKinnon, J. A., T. M. S. Johnston, and R. Pinkel (2008), Strong transport and mixing of deep water through the Southwest Indian Ridge,
Nat. Geosci., 1(11), 755–758, doi:10.1038/ngeo340.
Meehl, G. A., J. M. Arblaster, J. T. Fasullo, A. Hu, and K. E. Trenberth (2011), Model-based evidence of deep-ocean heat uptake during
surface-temperature hiatus periods, Nat. Clim. Change, 1(7), 360–364.
Melet, A., M. Nikurashin, C. Muller, S. Falahat, J. Nycander, P. G. Timko, B. K. Arbic, and J. A. Goff (2013), Internal tide generation by abyssal hills
using analytical theory, J. Geophys. Res. Oceans, 118, 6303–6318, doi:10.1002/2013JC009212.
Munk, W., and C. Wunsch (1998), Abyssal recipes II: Energetics of tidal and wind mixing, Deep Sea Res., Part I, 45(12), 1977–2010.
Niedzielski, T., Å. Høines, M. A. Shields, T. D. Linley, and I. G. Priede (2013), A multi-scale investigation into seafloor topography of
the northern Mid-Atlantic Ridge based on geographic information system analysis, Deep Sea Res., Part II, 98, 231–243, doi:10.1016/
j.dsr2.2013.10.006.
Nikurashin, M., and S. Legg (2011), A mechanism for local dissipation of internal tides generated at rough topography, J. Phys. Oceanogr.,
41(2), 378–395.
Nycander, J. (2005), Generation of internal waves in the deep ocean by tides, J. Geophys. Res., 110, C10028, doi: 10.1029/2004JC002487.
Pinkel, R., M. Buijsman, and J. Klymak (2012), Breaking topographic lee waves in a tidal channel in Luzon Strait, Oceanography, 25(2), 160–165.
Shapiro, G., and A. Zatsepin (1997), Gravity current down a steeply inclined slope in a rotating fluid, Ann. Geophys., 15, 366–374.
Shin, J., S. Dalziel, and P. Linden (2004), Gravity currents produced by lock exchange, J. Fluid Mech., 521, 1–34.
Simmons, H. L., S. R. Jayne, L. C. St Laurent, and A. J. Weaver (2004), Tidally driven mixing in a numerical model of the ocean general circulation,
Ocean Modell., 6(3–4), 245–263, doi:10.1016/s1463-5003(03)00011-8.
Smith, W. H. F., and D. T. Sandwell (1997), Global sea floor topography from satellite altimetry and ship depth soundings, Science, 277(5334),
1956–1962, doi:10.1126/science.277.5334.1956.
St Laurent, L., and C. Garrett (2002), The role of internal tides in mixing the deep ocean, J. Phys. Oceanogr., 32(10), 2882–2899.
St Laurent, L. C., and A. M. Thurnherr (2007), Intense mixing of lower thermocline water on the crest of the Mid-Atlantic Ridge, Nature,
448(7154), 680–683, doi:10.1038/nature06043.
Thorpe, S. (1977), Turbulence and mixing in a Scottish loch, Philos. Trans. R. Soc. London. Ser. A, Math. Phys. Sci., 286, 125–181.
Thurnherr, A. M., L. C. St Laurent, K. G. Speer, J. M. Toole, and J. R. Ledwell (2005), Mixing associated with sills in a canyon on the midocean
ridge flank, J. Phys. Oceanogr., 35(8), 1370–1381.
Van Haren, H. (2005), Details of stratification in a sloping bottom boundary layer of Great Meteor Seamount, Geophys. Res. Lett., 32, L07606,
doi:10.1029/2004GL022298.
Van Haren, H., R. Groenewegen, M. Laan, and B. Koster (2005), High sampling rate thermistor string observations at the slope of Great Meteor
Seamount, Ocean Sci., 1(1), 17–28.
Wunsch, C., and R. Ferrari (2004), Vertical mixing, energy and the general circulation of the oceans, Annu. Rev. Fluid Mech., 36, 281–314,
doi:10.1146/annurev.fluid.36.050802.122121.
Zilberman, N. V., J. M. Becker, M. A. Merrifield, and G. S. Carter (2009), Model estimates of M-2 internal tide generation over Mid-Atlantic Ridge
topography, J. Phys. Oceanogr., 39(10), 2635–2651, doi:10.1175/2008jpo4136.1.
©2015. The Authors.
491