OSOcean ScienceOSOcean Sci.1812-0792Copernicus PublicationsGöttingen, Germany10.5194/os-14-127-2018Observations of brine plumes below melting Arctic sea iceBrine plumesPetersonAlgot K.algot@uib.nohttps://orcid.org/0000-0003-0716-473XGeophysical Institute, University of Bergen, Bergen, NorwayBjerknes Centre for Climate Research, Bergen, NorwayAlgot K. Peterson (algot@uib.no)21February201814112713813April20178May201717January201822January2018This work is licensed under the Creative Commons Attribution 3.0 Unported License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/3.0/This article is available from https://os.copernicus.org/articles/14/127/2018/os-14-127-2018.htmlThe full text article is available as a PDF file from https://os.copernicus.org/articles/14/127/2018/os-14-127-2018.pdf
In sea ice, interconnected pockets and channels of brine are surrounded by
fresh ice. Over time, brine is lost by gravity drainage and flushing. The
timing of salt release and its interaction with the underlying water can
impact subsequent sea ice melt. Turbulence measurements 1 m below melting
sea ice north of Svalbard reveal anticorrelated heat and salt fluxes. From
the observations, 131 salty plumes descending from the warm sea ice are
identified, confirming previous observations from a Svalbard fjord. The
plumes are likely triggered by oceanic heat through bottom melt. Calculated
over a composite plume, oceanic heat and salt fluxes during the plumes
account for 6 and 9 % of the total fluxes, respectively, while only lasting
in total 0.5 % of the time. The observed salt flux accumulates to
7.6 kg m-2, indicating nearly full desalination of the ice. Bulk
salinity reduction between two nearby ice cores agrees with accumulated salt
fluxes to within a factor of 2. The increasing fraction of younger, more
saline ice in the Arctic suggests an increase in desalination processes with
the transition to the “new Arctic”.
Introduction
In the Arctic Ocean, sea ice is an effective barrier for exchange between the
ocean and atmosphere. The presence of sea ice, however, depends on a
delicate balance between the atmospheric and oceanic heat fluxes. The
inflowing Atlantic water contains enough heat to melt the Arctic sea ice in a
few years , and a small change in oceanic heat flux can
have huge implications for the heat balance at the interface. Understanding
the processes that control vertical heat fluxes under the sea ice is
important to understand the response of sea ice to a changing climate
. The interplay between heat and salt exchange at the
ice–ocean interface can work to enhance or reduce sea ice melt in the Arctic
Ocean .
While sea ice in bulk is a source of fresh water to the upper ocean, the sea
ice consists of fresh ice surrounding pockets of liquid brine connected
through a network of channels and capillaries . The brine
remains at its salinity-determined freezing point in thermal equilibrium
with the surrounding ice, and brine salinity and volume adjust to
temperature changes by growing or melting fresh ice.
Over time, salt is lost from the sea ice. The timing of the salt release and how
the salt is distributed in the water column is important in the evolution of
the Arctic mixed layer. The main desalination processes of sea ice are
gravity drainage and flushing of surface meltwater and melt ponds
. While melt ponds are present only in advanced stages of
melt, gravity drainage occurs throughout the seasons. Ice permeability is a
controlling factor for gravity drainage, increasing with temperature as the
ice warms . When sea ice warms to within a certain critical
temperature range, full-depth brine convection and desalination can occur,
even before the onset of melt . Furthermore, gravity
drainage has been successfully modeled using a 1-D sea ice model and can be
triggered both by atmospheric heat and bottom melt from oceanic heat .
Despite the theoretical understanding and successful modeling of spring-time
brine convection, observations are sparse. Brine drainage in response to
atmospheric warming may have been the cause of observed salinity anomalies
below sea ice in Storfjorden . Still, the main evidence so
far has been observations of saline plumes descending from warming landfast
sea ice in a Svalbard fjord . It has been hypothesized that
this form of desalination can occur on drifting Arctic sea ice, but so far
this has remained unverified. The existence of such plumes can be important
to the desalination of sea ice and the subsequent distribution of salinity in the
upper water column, and they could thus affect the otherwise strong surface
stratification typical below melting ice.
The first observations of brine plumes released from drifting sea ice in the
Arctic Ocean are presented here. The observations were collected in June 2015
in the marginal ice zone (MIZ) north of Svalbard (Fig. ). The data are a
subset of a previously reported under-ice turbulence data set
and part of the Norwegian Young Sea
Ice Cruise N-ICE2015;.
Data and methodsTurbulence instrument cluster
Under-ice turbulence measurements were made using a turbulence instrument
cluster (TIC) deployed 1 m below the ice undersurface, relying on
eddy covariance to calculate turbulent fluxes of momentum and scalars from
point measurements of temperature, salinity, and currents. The cluster is
fixed on a mast deployed through a hole in the sea ice and suspended on
a wire, which allows for the adjustment of the instrument depth. The concept is
well proven, and processing follows previous studies
. A detailed description of the setup is given in
and briefly summarized below. Horizontal and vertical
currents are rotated into the mean current direction (u) such that the
cross-stream (v) and vertical (w) current averages zero for a given
15 min segment. The data gaps visible in Fig. are
due to two corrupt data files.
Heat flux is calculated from the covariance of temperature and vertical velocity,
FH=ρwcp〈w′T′〉,
where ρw is the water density and cp is the specific heat capacity
of the water; angled brackets indicate a temporal mean, and primes indicate
detrended values (fluctuations about a 15 min mean value). The heat flux
is positive when warmer water is brought upward, and cold downward.
Similarly, salinity flux is calculated as
FS=〈w′S′〉,
where FS is positive when more saline water is brought upward, and fresher
water moves down. Accumulated salt flux (units kg m-2) is calculated by
adding up 15 min salt fluxes (m s-1) multiplied by the segment's
duration (s) using salinity in kg m-3.
Overview of the study region north of Svalbard, showing the drift
track of the ice camp studied here, with its starting point marked by a
cross. The positions of the Van Mijenfjorden study and
Whaler's Bay are shown for
reference.
Turbulent fluxes of (a) heat and (b) salt and
(c) friction velocity are shown as 15 min data points (dots) and
3 h bin averages (diamonds). In (a), sea ice thickness is shown
from manual measurements in the TIC hole (circles), hot wires (line), and two
ice cores (stars). (b) Total salt content of the ice cores (stars)
and measured cumulative salt flux, Stot, is given in kg m-2
(gray). Identified plumes are indicated by blue triangles. In (c),
the sea ice drift speed (thin) and along-stream current (TIC, thick) are
shown (gray), and the timing of a passing storm is indicated by the green line
defined by.
The covariance of horizontal to vertical velocity gives the components of
Reynold's stress, presented here as friction velocity,
u*=τ=〈u′w′〉2+〈v′w′〉21/4,
where τ is the kinematic Reynolds stress magnitude.
The TIC data and the derived fluxes have been subjected to an extensive
quality control, which is described in full in . The
systematic approach is taken to ensure the validity of Taylor's hypothesis,
which is crucial to the turbulent flux calculations. Taylor's hypothesis
assumes that an eddy is essentially unchanged as it passes the measurement
volume, allowing temporal measurements to be translated into spatial
measurements. Each 15 min segment is split in 1 min half-overlapping
subsegments, for which mean and root mean square values are calculated. This
is compared to artificial Gaussian data and used to identify variability
in the flow that violates Taylor's hypothesis, such as trends, rapid change
in current direction, and swell. Segments that do not meet the criteria
indicate unsteady flow and are excluded from the analysis. For the data
presented here, 85 out of 612 segments (14 %) were rejected in quality
control. Additional details on processing and data considerations are
discussed in Sect. .
Auxiliary data
The TIC data are supplemented by atmospheric data from a 10 m tall weather
mast and navigational data from the research
vessel Lance, which was anchored to the same ice floe during the drifts
approximately 300–400 m away.
Environmental data from the upper ocean are obtained from profiles of
temperature and salinity made using a microstructure sonde MSS,.
The profiles were typically collected in sets of three casts
repeated three times daily. Casts were made through a hydrohole about 50 m
from the TIC site. Data were validated against the shipborne CTD
(conductivity, temperature, depth) and corrected for sensor drift. The data
were analyzed using the Thermodynamic Equation of Seawater 2010
TEOS-10,, and conservative temperature (Θ) and
absolute salinity (SA) are used throughout.
Ice cores were sampled throughout the campaign for different ice types and
sampling variables . Two colocated ice cores with both
temperature and salinity measurements were collected on the same ice floe as
the flux measurements and are used in this study. Brine volume is calculated
as Φ=Sbu/Sbr, where Sbu is the bulk salinity, and brine
salinity is calculated using the linear relation Sbr=-Tice/0.05411,
which is adequate for warm ice .
In addition to the total height of the two ice cores, sea ice thickness is
measured manually in the TIC hydrohole and in a grid of hot wires (Fig. ).
The manual measurements were read from a ruler (Amelie Meyer, personal
communication, January 2017). Due to large but unknown uncertainty,
the measurements in Fig. a are arbitrarily assigned
±15 cm error bars. A set of four hot wires were set up in an area of
deformed sea ice, initially nearly 2 m thick . The error
bars in Fig. a are the standard deviation of the wires.
Because of the spatial variability and uncertainties of the different
measurements, all ice thickness data should be interpreted with care.
Environmental setting
Observations were made from a drifting ice floe in the MIZ between 10 and 19 June
(Figs. and ). The drift took place over
the Yermak Plateau, where a branch of the warm West Spitsbergen Current flows
across the plateau . Over 9 days, the floe drifted 185 km,
with an average drift speed of 23 cm s-1, while water depths shoaled
from about 2000 m to less than 1000 m over the Yermak Plateau. The floe had
an approximate diameter of 1200 m and likely consisted of only first-year
ice . The TIC mast was deployed approximately 250 m from
the floe edge. The ice drift was mostly parallel to the ice edge (Fig. ).
Ocean and sea ice conditions over the course of the drift.
(a) Drift track between 10 and 19 June (black) with daily ticks
(crosses). Conservative temperature (colors) data are derived from a
combination of vertical MSS profiles in the upper 30 m and TIC time series
measurements from 1 and 5 m. The ice edge (50 % concentration) is shown
for 12 and 18 June. Water depth is indicated in shading, with yellow isolines
at 1000 and 2000 m. Triangles mark the location of (b) four
profiles of stratification (N2) in the upper 50 m.
Temperature at 1 m below the ice averaged ΔT= 0.6 ∘C above
freezing and was lowest on 11 June (ΔT= 0.1 ∘C) and highest
(1.6 ∘C) during the storm on 13 June. Atlantic water flows along the
topographic slope and is often found at depths shallower
than 30 m (defined as T> 0 ∘C, Fig. ). Toward the end
of the drift a warm intrusion is also observed at 5 to 10 m depth.
Stratification (Fig. b) in the upper 35 m varies
significantly over the drift in and out of warmer waters. The mixed layer
depth gradually changes from quite deep (> 30 m) in the beginning of the
drift to non-existent at the end, varying with drift to and from areas where
warm Atlantic water flows closer to the surface. First, there is a transition
from waters of weak stratification (11 June) to gradually stronger surface
stratification. On 12 June, the top of the pycnocline is about 20 m, reaching
27 m on 14 June. Towards the end of the drift, there is strong stratification
continuously up to the surface, and there is no mixed layer present on 18 June.
Although sea ice thickness measurements are coarse, significant melt is
evident over the drift (Fig. a). The measurements in the
TIC hydrohole indicate a reduction from ∼ 100 to ∼ 40 cm between
12 and 18 June. Less melt was seen from the ice cores, with a reduction from
109 to 89 cm between 13 and 17 June. Hot wires measured a decrease from
174 to 87 cm over the measurements, although with a very large difference
between sensors (variance of up to 46 cm). The ice around the hydrohole is
likely melting faster compared to some distance away, and a representative
sea ice reduction is likely somewhere between hydrohole and ice core values.
Still, by the end of the measurements on 19 June, the ice was only a few
decimeters thick, and the floe was disintegrating.
Vertical profiles of sea ice (a) temperature,
(b) bulk salinity, and (c) brine volume fraction. The ice
cores are sampled about 100 m from the measurement site on 13 and 17 June.
Average temperatures are -1.9 and -1.4 ∘C, and bulk salinities
are 6.4 and 4.8 for 13 and 17 June, respectively. The typical 5 % threshold
required for gravity drainage is indicated (c,
dotted line).
From an ice coring site located approximately 100 m from the measurement site
but on the same ice floe, two ice cores sampled on 13 and 17 June
give some insight (Fig. ). The ice core on 13 June
shows 109 cm thick ice, with a 30 cm snow layer on top. The ice is rather
warm, with a minimum temperature of -2.5 ∘C in the interior of the ice,
increasing towards the surface (-1.3 ∘C) and the ice–ocean interface
(-1.7 ∘C). The C-shaped temperature profile is indicative of a
gradual warming from above. This is confirmed by atmospheric measurements
reporting temperate conditions throughout the measurements on the floe, with
temperatures ranging from -2 to +2 ∘C at 10 m of height between 7 and 20 June
. The snow layer was thick (∼ 30 cm), slowing heat
exchange with the ice . Comparison of the two ice cores
reveals a decrease in bulk salinity from 6.4 to 4.8 in 4 days and
a decrease in thickness of 20 cm, together causing a decrease in salt content
of 2.8 kg m-2 (calculated by multiplying bulk salinity with ice thickness).
Results
Eddy covariance measurements from 1 m below the ice undersurface reveal
anticorrelated turbulent heat and salt fluxes (Fig. ,
r=-0.94) at a time of rapid bottom melt. Oceanic heat fluxes are directed
towards the ice and reach several hundred W m-2 in response to a
passing storm. Salt fluxes are directed down from the ice, exceeding
-10-4 m s-1. The downward flux of salt is typical of freezing
conditions, such as that observed in refreezing leads in the pack ice north
of Alaska . During melting conditions, heat and salt fluxes
are more typically both positive, as fresh meltwater is fluxed downward and
is replaced by warmer water from below.
At the surface we generally find cooler, fresher water than below (Fig. ),
which is consistent with observed melting at the surface. The
negative salt flux can thus not be caused by the entrainment of saline water from
below. The observed heat and salt fluxes implies relatively cool, saline
water above warmer, fresher water, which is an inherently unstable
configuration that cannot be sustained over time. Unstable conditions can
occur during a frontal passage, during which the observation point (ice floe) drifts
from cool and saline water into an area of warmer, fresher water
. When the floe drifts into recently ice-free
waters, freshened from sea ice melt and warmed by the sun, cool water moving
with the ice floe could be dragged over warm water, setting up an instability
with appropriate gradients. The floe drifts over recently ice-free waters on
two occasions, and for shorter periods such overturning might be expected,
most notably in the last hours on 12 June concurrent with the decreasing
flux magnitudes. However, the negative relationship between heat and salt
fluxes is sustained over several days, during both increasing and decreasing
temperatures, signaling a process that is continually feeding the
instability. Negative correlation between the fluxes is consistent throughout
the measurements. The turbulent heat flux is a likely forcing agent, as both
fluxes increase with drift speed (Fig. c) and upper
ocean temperature (Fig. a).
Brine released from warm sea ice is a possible explanation, as a source of
cold, salty water sinking from the ice into the first meter of the ocean is
consistent with negative salt flux and positive heat flux. Resemblance to the
observations in the fjord study by inspired the search for
an inferred mean plume structure. Events are identified in a similar manner,
requiring at least five consecutive points at which w′< 0, S′> 1 × 10-5
and the salt flux magnitude, |w′S′|, exceeds 10-4 m s-1 or at
least 5 times the root mean square value over the 15 min segment. A
60 s window centered on the peak w′S′ value is used to construct a
mean plume ensemble. For each iteration, the 15 min window is moved 5 min
in order to also detect plumes otherwise falling on the edge of a
window. Duplicate events are removed, leaving 131 identified plumes for the
ensemble average, as shown in Fig. . Averaging is done using a
bootstrap calculation , which resamples the data 1000 times
to obtain an estimate of the average value occurring by chance. The mean
plume and its 95 % confidence interval from bootstrap calculations are shown
in Fig. . The shading represents percentiles of the data as
a display of the variability between plumes.
Composite of 131 plumes identified using the peak in
〈w′S′〉 and presented in a 60 s window. Spread in the data is
shown as percentiles (shading) overlain by the mean (white) and its 95 %
confidence interval from bootstrap calculations (dashed red). Variables are
fluctuations of (a) horizontal and (b) vertical velocity,
(c) salinity, (d) temperature, and turbulent fluxes of
(e) salt and (f) heat. The results from
are shown for reference (black dotted lines).
The inferred plume is approximately symmetric in time about its peak.
Anomalies in temperature and salinity gradually increase toward their peak
values over about 10–15 s before they decrease again at the same pace.
Vertical velocity perturbations, and thus also the fluxes, increase more
abruptly, reaching a peak of 2–6 cm s-1 in about 7 s before returning
to near zero. Horizontal velocity typically retards by around 5 cm s-1
during the plumes. Temperature and salinity anomalies deviate somewhat from
symmetry, averaging positive (2.1 × 10-3∘C and
4.0 × 10-3) before the plume and approaching zero after. There is,
however, considerable variation between individual events, and these are just
characteristics of the mean structure. Individual plumes typically have
sharper interfaces, and the smooth transitions in Fig. are
partly due to averaging. Salt and heat fluxes averaged over the 14 s
surrounding the inferred plume peak are FS=-3.7 × 10-4 m s-1
and FH= 1058 W m-2, respectively. The values observed here are up to
1 order of magnitude greater than those found by (see Table ).
Statistics of fluctuations and turbulent fluxes in the present study
over the Yermak Plateau in comparison with other Arctic studies of turbulent
heat and salt fluxes.
Location and studyStatisticw′S′T′FSFH(cm s-1)(10-3)(m K)(10-4 m s-1)(W m-2)Open leads 1 h mean-0.155.9Van Mijenfjorden Plume peak-1.610-3.7-1.5215Whaler's Bay Mean0.19268Overall mean-0.1975Yermak PlateauPlume mean-2.018-11-3.71058Plume peak-4.223-15.8-8.72465
The impact of drift velocity on the plume observations is investigated in
Fig. . Drift speed does not relate linearly with the
maximal heat flux in the plumes. In fact, many of the most intense plumes
observed (highest FH) are during weak or moderate current speed. The peak
in vertical velocity increases with increasing current speed (Fig. b).
This does not imply that plumes are stronger during fast
drift, but rather that wmax is dominated by increasingly large turbulent
eddies. Stronger turbulence causes more efficient entrainment of ambient
water into the plumes, leading to a lower peak in turbulent heat flux. Drift
speed also relates to the deviation in temperature from freezing,
ΔT=T-Tf, calculated from mean SA and Φ over the 14 s
surrounding the peak in vertical velocity. For low drift speed, many of the
plume observations carry water that is supercooled relative to the ambient.
The supercooling is caused by the lower salinity-determined equilibrium
temperature of the brine within the sea ice. Supercooling decreases with
drift speed and is not observed for plumes in which the mean current exceeds
∼ 25 cm s-1, which is consistent with stronger mixing of the plumes during
high drift speed. Plumes associated with high maximum heat fluxes are more
often supercooled than not.
Mean horizontal current (Um) vs. instantaneous
(a) maximum heat flux and (b) minimum vertical current
speed. Circles are color coded for temperature above freezing
(ΔT=T-Tf), calculated using mean absolute
salinity and conservative temperature . Mean values are
calculated over the 14 s surrounding the peak vertical current speed of each
identified plume.
The accumulated salt release during the flux observations was
7.6 kg m-2 summed over available measurements between 11 and 19 June
(Fig. b). This is equivalent to a salinity decrease of 5
for 1.5 m thick ice. The salt flux observed here is approximately equivalent
to the total salt content of the 13 June ice core. About half of the salt
flux was observed before the ice core was sampled, so desalination had
already taken place before coring. The salt flux observed after the time of
coring accounts for 57 % of the total salt content of the ice core on
13 June. The salt flux measured between the time of the two cores is
2.8 kg m-2, the same amount as the change in salt content of the two ice
cores. However, gaps in the time series point to a discrepancy between ice
cores and observed salt flux. Assuming the salt flux during the measurement
gaps was equal to the mean of available measurements, the accumulated salt
flux between the two ice cores is approximately twice the observed reduction
in the ice cores. The discrepancy might be linked to spatial inhomogeneity in
ice composition and melt rates, variability between individual ice cores, or
measurement errors (Sect. ). Agreement between measured
fluxes and salinity in ice cores within a factor of 2 supports the fact that the
salt flux can originate in brine release from the sea ice.
Calculated over the 14 s surrounding the peak salinity flux in the mean plume
structure, the 131 identified events account for 0.7 m s-1, or 9 %, of
the observed total salt release within a duration of 31 min (0.5 % of
the time), illustrating the intensity of the events. The heat flux averaged
over the composite plume is 1058 W m-2, and the plumes account for 6 %
(4.7 W m-2) of the average observed heat flux between noon on 11 June
and
the end of measurements on 19 June. However, the plumes can additionally
cause mixing of the surface layers, which could counteract the stabilizing
effects of bottom melt. The overall effect of the plumes on heat fluxes is
thus difficult to quantify. Upper ocean hydrography profiles
Fig. ; do not provide conclusive evidence, as
advection and mixed layer deepening from wind forcing obscures any effect from the plumes.
Percolation or flushing of melt ponds could influence the measurements.
Although the first melt pond was noted on 9 June, they remained at a very
early stage throughout the measurement period reported here. The pond
fraction reached an estimated 10 % coverage. Mostly, ponds had formed at
deformation areas where freeboard was negative and were thus flooded with
seawater rather than actual being melt ponds (A. Rösel, personal communication,
14 January 2017). Salinity measurements from three melt ponds revealed an absolute
salinity of 20–29 g kg-1. The ice core from 13 June
(Fig. ) had a 2 cm negative freeboard, and the deepest
snow layer had a salinity of 4.3. Based on this and noting the high
permeability of the ice (high liquid fraction; Fig. c),
percolation may have played a role in the desalination process, but is not
pursued further here.
The combined heat flux from above and below finally melted the sea ice.
Substantial melt is also evident from the different ice thickness
measurements (Fig. a). At the end of the flux
measurements there were only a few decimeters of ice left, and the floe
disintegrated as the instruments were recovered on 19 June. Over the course
of the measurements, bottom melt caused an overall reduction in salinity
measured at 1 m by approximately 1.
Discussion
The observations of saline plumes presented here extend the findings of
and are the first observations of such plumes from
drifting Arctic sea ice. While the structure is similar to the observations
from Van Mijenfjorden , which were made with the same
instrumentation on landfast ice, the magnitudes observed here are much
greater, with peak values of salt and heat fluxes in the average plume of
FS=-8.7 × 10-4 m s-1 and FH= 2465 W m-2 (see
Table ). While the measurements by were made during
little (or no) ice melt, the present observations were made during severe
melting, which may be the primary difference between the two studies. The
fjord study concluded that oceanic heat from the tidal inflow likely
triggered brine release from the temperate ice .
During melting conditions, desalination can happen by gravity drainage or
flushing . Flushing can occur when there is an overhead
pressure from meltwater at the surface. The negative freeboard in the ice
core and saline melt ponds indicate that there was in fact overhead seawater
at the surface, which may have caused or at least increased desalination.
Gravity drainage occurs when the buoyancy of the brine exceeds the
dissipative effects of thermal diffusion and viscosity within the sea ice.
Atmospheric cooling in winter causes higher brine salinity and thus density
in the upper part of the ice column. This makes the brine unstable, and
convection within the ice takes place when the ice is sufficiently permeable
. As the ice warms in spring, permeability increases as
illustrated in Fig. . The brine fraction in the ice
cores exceeds 10 % throughout the ice (Fig. c), both well
above the typical 5 % threshold required for gravity drainage
. The instability needed for full-depth brine convection in
the ice can be triggered by atmospheric or oceanic heat or a
combination of the two . The temperature at the ice–ocean
interface is always at its salinity-determined freezing point, and warming
from below can only be caused by freshening at the interface by ice melt.
To the left is an early spring situation in which the upper ocean is
near freezing, and temperature in the ice is still below the critical
temperature Tc, which must be exceeded for gravity drainage to
occur. When the atmosphere warms the ice, permeability of the sea ice
increases, and gravity drainage can occur. The brine plumes are triggered by
meltwater below the ice, by directly exposing brine pockets, or by elevating
the freezing point temperature Tf at the interface. The TIC mast
is shown for reference, with measurement volume at 1 m below the
ice.
Ice melts at the ocean interface when the heat supplied by the ocean exceeds
the conductive heat flux in the ice. The high oceanic heat fluxes are caused
by a combination of the passing storm, the presence of warm Atlantic water near
the surface, and the high mobility of the sea ice driving mixing
; this accounts for much of the observed melt.
As the interface freshens, the interface's salinity-determined temperature
Tf(S) increases (Fig. ). Tf remains high as
long as the fresh water is allowed to remain at the interface, or enough new
meltwater is supplied. Additionally, fresh water supplied by melting at the
ice–ocean boundary increases the density deficit between the lower part of
the ice and of the brine, which may increase the potential for gravity drainage.
The various ice thickness data (Fig. a) show that
significant melt was indeed occurring, up to as much as 25 cm day-1
during the observation period. Furthermore, the present observations show
positive temperature anomalies prior to the plume events (Fig. c).
The positive temperature anomaly before the plumes, the
sustained positive heat fluxes (Fig. a), and the rapid melt
suggest that oceanic heat plays a key role through ice melt, as required for
triggering repeated convection events.
The difference in salt content between the two ice cores is mostly
(∼ 75 %) due to a reduction in ice thickness. The rapid ice melt is thus
the cause of most of the salt release and likely the reason for the large
difference between values observed here and those of ,
in which little or no ice melt took place. In addition to gravity drainage or
flushing of brine, the brine pockets become directly exposed as melt
progresses and sink past the measurement volume. Considering that most of the ice
volume is lost during the measurements, it is not surprising if most of the
original brine content in the ice is lost. It is, however, surprising that
measurements at 1 m below the ice are of the same order of magnitude as the
total desalination, which calls for an investigation of possible measurement
errors or biases (Sect. ).
When sea ice melts, it contributes to a net freshening of the upper ocean,
since the bulk salinity is about 5 (Fig. ). Over the
course of the drift, a freshening of the surface layer is observed, while the
turbulence measurements at 1 m show negative salt flux. This paradox warrants
some consideration of the structure of sea ice. Sea ice consists of
freshwater ice surrounding pockets of liquid high-salinity brine. When the
sea ice melts, the brine sinks through the surface layer due to its high
density, while the fresh meltwater stays at the surface. The fresh surface
water is gradually entrained into the mixed layer, but since the salt flux is
nearly always negative, this freshwater flux likely occurs on timescales
longer than the 15 min segments used here. Why fresh water at the surface
is not immediately mixed down, even during quite strong mixing events, is not
entirely clear. The ice floe consists of first-year ice ,
but the floe was deformed through several storms. This is evident, for example, from
the hot wire measurements, which were made in an area of deformed ice. A
rough undersurface of the sea ice leads to a thicker layer in which molecular
viscosity and diffusivity are important. This “transitional sublayer” is
usually taken as 1/30 of the scale of the roughness elements and is on the
scale of a few centimeters .
Since the salt flux observed within the 131 identified plumes only accounts
for 9 % of the total salt flux, most of the salt flux takes place outside
these plumes. Many more plumes are likely present nearby, but do not reach
or cross the measurement volume. Such plumes would bring higher-salinity
water somewhere above the TIC, rather than being immediately mixed in with
the fresh water at the ice–ocean interface. Subsequent mixing would be
observed as a negative salt flux, although not identified as a plume. This
may be the reason why salt fluxes are consistently negative, even though
direct plume observations are more sporadic.
Brine released from sea ice would initially be at its salinity-determined
freezing point in balance with the surrounding ice. As the brine descends
from the ice, it may thus be supercooled relative to its surroundings. When
the horizontal velocity (and u*) is greater, shear mixes and dilutes
the released salt plumes more than during calm conditions. This is consistent
with the observation of less supercooling with higher mean current, as seen
in Fig. . Maximum vertical velocity in the plumes
coincided with higher drift speed, which is likely caused by stronger
turbulent eddies. The individual plumes do not grow more intense with
stronger turbulence, but are more efficiently mixed into the surroundings.
The observations of brine plumes raise interesting questions concerning the
conditions in which they occur and, importantly, why they have not been
observed before. Few studies have reported measurements of turbulent salt
fluxes in the Arctic Ocean, and the season of the observations may be of the
essence. In the preceding ice camp (Floe 3) in the N-ICE campaign, the salt
fluxes were below the sensor accuracy level and could not be analyzed for
correlation with heat flux. The study by is relevant for
comparison because it was set in roughly the same place, Whaler's Bay, in
April 2003. The differences between this study and may
give some clues to the matter. They deployed the TIC in a refrozen lead
surrounded by ridged multi-year sea ice. An ice core revealed a linear
temperature gradient of -21.7 K m-1, meaning that the ice was not above
the critical 5 % threshold for gravity drainage. The cold ice column may
explain why brine was prevented from leaving the sea ice in plumes, despite
rapid melt. When brine is released slowly as melt progresses, it is more
likely to be mixed in with meltwater than to descend in plumes.
Error sources and biases
Salinity is calculated from a Sea-Bird Electronics SBE 4 conductivity cell. The
dependence of salinity on both conductivity and temperature can introduce
spurious salt fluxes because of a difference in response time between the
temperature sensor and the conductivity cell. The standard SBE 4 was chosen
for flux calculations rather than the SBE 7 micro-conductivity sensor because
the SBE 7 reported suspicious values for part of the record.
made a comparison of a ducted conductivity cell (SBE 4)
with a fast-response micro-conductivity sensor (SBE 7) and showed that most
of the covariance occurred at lower frequencies. About 75 % of the salinity
flux was resolved by the SBE 4. The present observations are obtained during
moderate to strong forcing (5–35 cm s-1 drift speed), which improves
the response time of the conductivity cell. It is advisable to interpret the
observed fluxes with this uncertainty in mind, but note that the fluxes from
the ducted conductivity are more likely an underestimate than an overestimate.
Considering the possibility of a baroclinic signal from the edge of the ice
floe contaminating the measurements, the vertical modal structure is
calculated from the profiles of buoyancy frequency shown in Fig. b.
The phase velocity of the first baroclinic vertical mode is
0.25–0.43 m s-1 for the four profiles. Taking the closest distance to
the floe edge of ∼ 200 m, this implies a timescale for a signal
originating at the floe edge of around 10 min. This is comparable to the
segment length used for flux calculations (15 min), which could violate the
validity of Taylor's hypothesis here. However, the systematic quality control
described in Sect. was designed to identify any violation
of Taylor's hypothesis and would thus have been excluded from the analysis.
Increased buoyancy frequency during the summer drift could affect the flux
measurements. While the typical buoyancy period was about 1 h for the most of
the drift (January through May), periods around 10 min or less were seen in
June. Oscillations with periods on the order of the 15 min segment length
could affect turbulent fluxes. However, recalculating the data set using
5 min segments revealed no significant differences, and
concluded that the systematic quality control had already flagged any
contaminated segments.
The hydrohole, through which the turbulence mast is deployed, can be
suspected to affect measurements, and lateral heating may have caused faster
melt in some radius around the hole. Still, the horizontal component of the
flow is larger than the vertical, and observations made at the TIC represent
conditions at the ice interface some distance away. Taking a vertical
velocity anomaly of 2–5 cm s-1 and the mean horizontal component of
ΔU∼ 10 cm s-1 (difference between drift velocity and current
measured at 1 m; Fig. ), a plume signal moves 2–5 m
in the horizontal over the 1 m vertical distance from the ice–ocean
interface. The swiftest vertical speeds are also typically associated with
large horizontal speed (Fig. ). This indicates that even
for the large vertical speed seen in the plumes, influence from processes
around the hydrohole is typically not expected.
The exact distance between the TIC measurement volume and the ice
undersurface may be important for the absolute values observed, as one would
expect plumes to gradually expand in size but weaken in terms of buoyancy
anomaly with distance from the ice. The manual measurements of ice thickness
are accompanied by adjustments of the instrument depth. After each ice
thickness measurement, the instrument was elevated to account for the ice
melt. Interpreted from notes on these adjustments, the measurement volume was
always at the correct depth within the range of the measurement uncertainty.
In Fig. a, the uncertainty is arbitrarily set to ±15 cm.
Overall, salinity decreases by about 1 over the course of the drift, as
measured by the instrument at 1 m. At the same time, the accumulated salt
flux accounts for an increase in salinity of 1.7 if distributed over the
4.7 m average mixed layer depth . The apparent
inconsistency is caused by the separation in timescales. Over longer
timescales, fresh water from ice melt is fluxed downwards, but is not
apparent in the turbulence record because each 15 min segment is detrended
before fluxes are calculated. The comparison between accumulated salt flux
and the salt contents of the ice core indicates that most of the brine in the
ice convects down past the surface layer, rather than blending in with the
fresh meltwater at the interface. Thus, it appears that brine plumes as observed
here affect the timing of the salt release, alter how the salt from sea ice
is distributed in the water column, and can be an important factor
influencing mixing during sea ice melt.
Concluding remarks
This study reports observations of inversely correlated heat and salt fluxes
below melting sea ice north of Svalbard. The evidence suggests that the
fluxes are caused by brine release from the sea ice as it melts, and a
significant fraction of the salt fluxes are seen descending past the
measurement volume in plumes. Desalination of sea ice similar to that
observed here likely occurs in the MIZ in spring in general, when enough
heat is present to trigger such events. The present desalination appears to
be forced by a combination of flushing, gravity drainage, and direct release
of salt through rapid melt caused by oceanic heat flux. Triggering by ocean
heat flux is less likely in the Arctic interior, both because there is
typically less heat available to be mixed up (e.g., less open water to be
warmed by insolation) and less mixing due to internal forces in the pack
ice. In the interior Arctic Ocean, sea ice is typically second- or multi-year
ice, which is thicker and less saline than first-year ice. Brine release in
the quantities reported here are thus more likely a MIZ phenomenon. With the
transition towards a more seasonal ice cover in the Arctic, the fraction of
first-year ice is increasing . First-year ice is more saline
, with an equivalently greater potential for brine
drainage. While this could indicate that saline first-year ice can melt
faster than fresher multi-year ice in otherwise similar conditions, the
transition to more FYI comes with increased freshwater runoff and increased
upper ocean stratification . Desalination appears to be a
significant process in sea ice melt, although small in comparison to frontal
processes and solar heating. Understanding desalination processes may still
be increasingly more important in the “new Arctic” and requires more
targeted field campaigns.
The following data sets are used in this study and are
publicly available at the Norwegian Polar Data Centre: turbulence instrument
cluster data , microstructure sonde (MSS) profiles
, meteorological data , ice thickness
from hot wires , and ice core data .
The author declares no conflict of interest.
Acknowledgements
The fieldwork has been supported by the Norwegian Polar Institute's Centre
for Ice, Climate and Ecosystems (ICE) through the N-ICE project. The author
would like to thank everyone involved in the fieldwork, and Amelie Meyer in
particular, for keeping the instruments going until the very end. Thanks to
Ilker Fer, Dirk Notz, Mats Granskog, Martin Vancoppenolle, Craig Stevens, and
the anonymous reviewers for useful discussions and input and to Miles McPhee
for making the instrumentation available. The author is supported by the
Research Council of Norway through project 229786, with additional
support from the Centre for Climate Dynamics at the Bjerknes Centre through
the grant BASIC: Boundary Layers in the Arctic Atmosphere, Seas and Ice Dynamics.
Edited by: Eric J. M. Delhez
Reviewed by: Craig Stevens and one anonymous referee
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