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Abstract
The paper presents
a new hypothesis along with supporting evidence that the Beaufort
Gyre plays a significant role in regulating the arctic climate
variability. We propose and demonstrate that the Beaufort Gyre
accumulates a significant amount of freshwater during one climate
regime (anticyclonic) and releases this water to the North Atlantic
during another climate regime (cyclonic). This hypothesis can
explain the origin of the salinity anomaly periodically found in
the North Atlantic as well as its role in the decadal variability
in the Arctic region.
Introduction
The present state of the Arctic Ocean and its influence on
the global climate system strongly depend on
the Arctic Ocean freshwater budget (Aagaard and Carmack, 1989,
hereinafter A&C)
because fluctuations in the freshwater
export can significantly influence the depth and volume of
deep water formation in the North Atlantic (NA)
and ultimately the strength of the global
thermohaline circulation.
The traditional approach for investigations of the freshwater
budget of the Arctic Ocean has been to perform a detailed
analysis of its major components including river runoff, the inflow
of waters from the Atlantic and Pacific Oceans,
the outflows through Fram Strait and the Canadian Archipelago, the
atmospheric moisture flux and the annual cycle of
ice formation and melt (see Lewis [2000]). Significantly
less attention has been paid to
the processes involved in the storage of FW in the Arctic Ocean
and its temporal variability. The regional differences in this
storage (e.g., in sea ice thickness and in ocean salinity) are
substantial ( A&C; Steele et al., 1996).
For instance, the Canadian Basin of the Arctic Ocean
contains about 45,000 km^3 of fresh water (calculated relative
to the salinity 34.80 by A&C).
This is 10-15 times larger
than the total annual river runoff to the Arctic Ocean,
and at least two times larger than the amount of fresh water (FW)
stored in the sea ice. A release of only 5% of this FW
is enough to cause a salinity anomaly in the North Atlantic
comparable in magnitude to the Great Salinity Anomaly of the 1970s.
The largest of the anomalies is
located in the Beaufort Gyre (BG), identified by a salinity
minimum at depths 5-400 m (Figure 1A-C (pdf)). This anomaly drives
the BG geostrophic circulation anticyclonically
(Figure 1D (pdf)). We propose that the freshwater
budget of the BG and the freshwater flux to the
NA depend significantly on the intensity of
this salinity anomaly and climatic conditions conducive to
the transport of FW from the BG to the NA.
This paper provides a step in
understanding the origin of this anomaly and the nature of its
variability. Characteristics and sources of the data for each figure are presented
in Table 1.
Beaufort Gyre as a flywheel
The origin of the salinity minimum in the
BG can be inferred by a comparison of the seasonal
change in wind and sea ice motion. Figure 2 (pdf) shows the wind and ice drift
patterns seasonally averaged
for the period 1979-1997. In winter (September-May), the wind (Figure 2A (pdf)) drives the ice
and ocean anticyclonically (Figure 2C (pdf)) and the ocean
accumulates potential energy through a deformation of the
salinity field (Ekman convergence and subsequent downwelling,
see Figure 1C (pdf)). The strength of the horizontal salinity
gradient and resultant geostrophic circulation
depend on the intensity and duration of the anticyclonic winds. During the
winter season the wind-driven and geostrophic currents coincide
to set up a strong anticyclonic ice rotation (Figure 2C (pdf)).
In summer (June-August), the wind is weaker or it may even be cyclonic
(Figure 2B (pdf)) but in the mean
the ice still rotates anticyclonically (Figure 2D (pdf)). An obvious
conclusion is that in summer the ocean geostrophic circulation prevails
and drives the ice against the wind motion. The
salinity anomaly and freshwater content (FC) in the BG (Figure 1B (pdf)) must decrease in
summer, because without wind support, the ocean
loses potential energy, i. e., Ekman pumping is reduced. During the
following winter the ocean again accumulates potential
energy. Hence, the climatic structure of the salinity and dynamic
height distribution remain rather persistent (not shown) although
exhibiting some seasonal and interannual variability.
When viewed on a seasonal scale, the BG salinity anomaly stabilizes the
circulation, remaining essentially anticyclonic throughout the year,
thus permitting the BG geostrophic circulation cell to serve as a flywheel
for the Arctic Ocean circulation.
Some modeling results confirming this mechanism are shown in Figure 3 (pdf).
An idealized situation has been tested using a 3-D Blumberg and Mellor
[1987] numerical model in a 2000x2000 km basin with 1500 m depth. The basin
was initially
horizontally uniform but vertically stratified, then it was forced for 9 months
by symmetric anticyclonic winds followed by 3 months
of cyclonic symmetric winds. The anticyclonic winds generate downwelling
in the central basin and upwelling along the boundaries (Figures 3A-B (pdf)).
The results after anticyclonic forcing only are similar to
the winter Arctic conditions, and the salinity structure in Figure 3B (pdf)
resembles that in Figure 1C (pdf).
The addition of cyclonic winds leads to upwelling in the central basin
and downwelling along
the boundaries and
to a reduction in the anomaly in the salinity field generated by anticyclonic winds.
The distribution of salinity and currents after 3 months of
cyclonic wind forcing are shown in Figures 3C-D (pdf). The circulation pattern in Figure 3C (pdf) is similar to the ice drift
pattern in Figure 2D (pdf), i.e., it is still anticyclonic
but is weaker than in winter. The salinity distribution in Figure 3D (pdf) resembles the summer
salinity distribution in Figure 1B (pdf) when the cyclonic wind forcing leads to the release of FW from
deep to upper layers.
The seasonal variability
of FW content in the central part of the basin is about 10% (not shown).
This seasonal mechanism of freshwater accumulation and release
is extended to the decadal time scale in the next section.
Table 1. Characteristics of Data
| Parameter |
Period
Span |
Reference
Period |
Figure |
Data
Source |
| Salinity (S) |
1970-1979 |
Winter |
Fig. 1A 1C |
EWG(a) |
 |
| S |
1970-1979 |
Summer |
Fig. 1B |
EWG(a) |
|
| T-S |
1970-1979 |
Jan.-Dec. |
Fig. 1D |
EWG(a) |
|
| S |
1973-1979 |
Mar.-May |
Fig. 4A |
AARI(b) |
|
| Buoy Drift |
1978-1997 |
Jun.-Aug. |
Fig. 2D |
IABP(c) |
|
| Buoy Drift |
1978-1997 |
Sept.-May |
Fig. 2C |
IABP(c) |
|
| SLP |
1978-1997 |
Sept.-May |
Fig. 2A |
NCAR(d) |
|
| SLP |
1978-1997 |
Jun.-Aug. |
Fig. 2B |
NCAR(d) |
|
| SIC |
1978-1997 |
Sept.-May |
Fig. 5A,5B |
NSIDC(e) |
|
a Environmental Working Group Atlas [1997,1998]
b Arctic and Antarctic Research Institute (pers. com).
c International Arctic Buoy Program
d National Center for Atmospheric Research
e National Snow and Sea Ice Data Center
Hypothesis
A hypothetical chain of
relationships among atmosphere, ice and ocean in the Arctic at the decadal
time scale
has been proposed by Mysak and Venegas [1998], Proshutinsky et al., [1999] (hereinafter P99) and others but it
is important to know what causes the variability.
In order to explain the relationship between the wind-driven and
geostrophic circulation and their
influence on the accumulation and release of FW we examine the interplay between the
atmosphere, ice and ocean in terms of the two circulation regimes identified
by Proshutinsky and Johnson [1997] (hereinafter P&J) and P99.
ACCR
During the anticyclonic circulation regime (ACCR), when high
atmospheric pressure prevails in the Arctic, the
Arctic Ocean accumulates FW through the increase of FW
volume in the BG (Ekman convergence and subsequent downwelling, see Fig. 1C) and
through the increase of ice thickness and area due to
enhanced ice growth (the Arctic is colder
during an ACCR than
a cyclonic circulation regime (CCR) as shown in P99). Ice is
additionally accumulated in the BG during an ACCR due to convergence
and ridging under anticyclonic wind forcing.
River
runoff is increased (trajectories of cyclones are shifted
toward land) (P&J; Johnson et al., 1999) and more FW
accumulates in the surface waters. When anticyclonic winds are
prevalent, the flow of Arctic waters towards Fram Strait is reduced
( P&J;
Trembley and Mysak, 1998).
Consequently, the ice and water flux from
the Arctic Ocean to the Greenland Sea and the transport
of Atlantic Water into the Arctic Ocean (as a compensation of outflow)
are weaker than usual. Deep convection in the Greenland Sea is then enhanced
because the vertical stratification is reduced (less FW in the surface
waters). This decoupling of the Greenland, Iceland, and Norwegian Seas (GIN Sea)
from the Arctic leads to their eventual warming.
Transition to CCR
All of the above processes lead (with some time lag) to an
increase in the gradient of dynamic height between
the BG and the NA.
The resultant geostrophic circulation increases as does the
outflow of FW and ice from the Arctic.
During
warming of the GIN Sea, the Icelandic
Low intensifies and moves to the north leading to an
intensification of the transport of
Atlantic waters into the Arctic Ocean. This increase in warm
water flux to higher latitudes enhances the penetration
of atmospheric cyclones into the Arctic, and ultimately
decreases the atmospheric pressure in
the Arctic. Warming of the Arctic establishes the CCR.
CCR
During the cyclonic circulation regime, when low
atmospheric pressure prevails in the Arctic (see table characterizing
different environmental features of CCR and ACCR in P99), the
Arctic Ocean releases FW to the NA through the passages in the Canadian
Archipelago and Fram Strait.
Warming in the Arctic during the CCR increases ice melting and
releases additional FW to the central basin.
The accumulation and storage of FW in the
BG is not favored by the CCR (even though the cyclonic regime
leads to increased ice
melt, the FW is not accumulated in the BG because
of Ekman divergence and upwelling causing a decrease
of freshwater volume in the BG), and hence more FW
is available for transport
to the NA. River runoff is lower during the CCR than during
the ACCR but precipitation over the ocean is increased and
hence more fresh water is available for immediate
release to the NA from sea ice and surface waters during the CCR.
The stronger surface winds of the CCR in the Fram Strait area
(P99) increase
the transport of thick ice, and hence FW, to GIN Sea.
At the peak of these processes,
when all of them coincide, we observe low salinity
anomalies in the GIN Sea.
Transition to ACCR
After several years of increased release of ice
and FW to the GIN Sea, the surface layer becomes cooler and fresher, and the
sea-ice extent increases in the Greenland Sea. Freshening associated
with melting of the increased ice volume and increased flux
of fresher surface waters leads to an increase in stratification and
a decrease in
the interaction between the deep ocean and the atmosphere;
deep water convection is consequently suppressed. After several years
the dynamic height gradient between the BG and the NA (and consequently the
geostrophic circulation) decreases, the Icelandic Low moves
to the south and the interactions between the GIN Sea
and the Arctic Ocean become weaker, reestablishing the
anticyclonic circulation regime.
It is important to note that in this sequence of processes
the accumulation and release of FW and ice plays a fundamental
role in the interaction between the Arctic Basin and the GIN
Sea.
Discussion and Conclusions
In order to support our hypothesis we have analyzed the variability of
the FC
in the BG (yellow box in Figs. 1 and 2)
using 1973-1979 March-May T-S surveys conducted
by the Arctic and Antarctic Research Institute (AARI, personal communication).
A time series
of the FC anomaly for this 7-year period is shown
in Figure 4A (pdf). Assuming that the FC in the BG depends on the intensity and
direction of
the wind-driven circulation, we correlated the FC with
the sea surface height gradient (SSHG) in the BG. This
gradient reflects the intensity of the anticyclonic/cyclonic wind-driven
circulation over the Arctic (see P&J and P99).
When the SSHG is positive, the ACCR prevails over the Arctic; when the SSHG is
negative, the CCR dominates. The time series of SSHG
anomaly for 1973-1979 (departure of SSHG from its mean for 1973-1979) is shown
in Figure 4A (pdf).
The correlation between the FC and SSHG
anomalies is 0.89. In order to expand this rather short time series to a longer
period, we employed a proportional relationship obtained by linear regression
(the FC anomaly = 2.*SSHG anomaly) to reconstruct the anomaly
of the FC in the BG for the period 1946-2001 (Figure 4B (pdf)).
The difference between the FC during ACCR
and CCR in the yellow box area
is about 10^4 km^3/year, which is about 3 times larger than the annual freshwater
input from river runoff estimated by A&C as 3300 km^3/year.
This suggests that the FW released from the BG during CCRs
can be significantly more important than that from
all other freshwater sources.
Another component of the FC is the volume of sea ice
in the Arctic Ocean. No direct observations are available but we can use some
results from modeling
studies. Figure 4B (pdf) shows the anomaly of sea ice
volume based upon the model studies
by Hilmer and Lemke, 2001.
This simulation reveals a pronounced
decadal variability of the sea ice volume which is in agreement
with the SSHG (except before 1970).
The total volume of sea ice in the three models is different but the sea ice
volume anomalies are quite similar.
The correlation between circulation regimes
and sea ice volume anomaly is excellent (but with some lag) after 1970.
In concept with our previous discussion, the volume of sea ice
increases during ACCRs and decreases
during CCRs. The disagreement noted for years prior to
1970 could be explained by model spin-up considerations.
Another confirmation of different rates of FW release from the Arctic Ocean
is the sea ice extent in the GIN Sea. Figure 5 (pdf) shows
the sea ice concentration (SIC) averaged for the ACCR and CCR years
since 1978 (see Table).
An enhanced development of sea ice
extending NW into the Greenland Sea
is noted during CCRs (Figure 5A (pdf)).
This provides indirect
evidence that deep convection is suppressed during
CCRs because of the large volume
of FW in the surface layer of the Greenland Sea.
At each location, if there were 3 or more
years that had ice over 20%, only those years were averaged. The 20% value
was used to keep out spurious
values over open water due to weather effects. Values of twenty percent or less really only occur in very narrow regions in the marginal ice zones.
One may wonder how the salinity anomaly in the BG
may change in response to
global warming and climate change. Recent observations
show that the climatically stable
ACCR, dominant during the 1980s,
has been replaced by
a CCR starting about 1989. As a
result, for most of the past decade the intensity of the Arctic
High has decreased and the
summer cyclonic circulation period (Figure 2B (pdf)) has commenced
earlier and lasted
longer than usual.
These conditions must necessarily lead to a salinity increase
in the deeper layers of the BG (upwelling in response to the cyclonic
forcing, similar to Figures 1B
and 3C-D (pdf)), a reduction in the speed of
the geostrophic current and to a decrease
of salinity in the upper layers of the Arctic Ocean. This latter is
due to suppression of Ekman
pumping reducing the transport of FW to the
deeper layers.
As a result, the FW stored in
the upper layers of the BG becomes available for output to the NA
through increased transport by the cyclonic
wind-driven circulation. Physical and geochemical data collected between 1989 and 1995
by McLaughlin et al. [2001] reveal that the the FC in the Canada Basin
has been significantly reduced
which confirms the reconstruction results of FC for 1990-1997 (Figure 4B (pdf)). Since 1997,
evidence suggests that
a new anticyclonic circulation regime may be developing ( P99) and we can expect
to observe an increase of the FC in the BG.
A substantial release of the BG
fresh water to the NA in response to changing climate
conditions can be
a source for a large scale salinity anomaly in the NA,
and consequently, a source for an abrupt global cooling
( A&C; Delworth et al., 1997).
The above perspectives lead
us to the conclusion
that it is extremely important to understand the structure of
the BG water properties, its currents, and their variability in space and time.
We encourage the creation of an observational and modeling
program to test the
hypothesis formulated above.
Acknowledgments
This research has been supported by a grant from NOAA.
It is contribution 10756 of the Woods Hole Oceanographic Institution.
References
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in the Arctic circulation, J. Geophys. Res., 94, 14,485--14,498, 1989.
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Arctic Ocean (CD-ROM). National Snow and Ice Data Center, Boulder, Colorado, 1997,1998.
Johnson, M. A., A. Y. Proshutinsky, and I. V. Polyakov, Atmospheric pattern
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McLaughlin, F., E. Carmack, R.W. MacDonald, A. J. Weaver and J. Smith,
The Canada Basin 1989-1995: Upstream events and far-field effects of the Barents Sea,
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Mysak, L. A., and S. A. Venegas, Decadal climate oscillations in the
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Steele, M., D. Thomas, D. Rothrock and S. Martin, A simple model study of the Arctic Ocean freshwater balance, 1979--1985, J. Geophys. Res., 101, 20,833--20,848, 1996.
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Figure Captions
Figure 1. (A) The salinity distribution at 25 m. (B),(C) Salinity distribution along dashed line in summer and winter. (D) Dynamic heights relative to 200 db and direction of geostrophic currents.
Figure 2. Winter (A) and summer (B) sea level pressure (SLP, hPa) and geostrophic wind. (C),(D) Seasonal sea ice drift.
Figure 3. Results of numerical experiments in the ideal basin. (A) Sea surface salinity (SSS) and surface currents. (B) Salinity section along dashed line. Both figures show results after 9 months of anticyclonic symmetric wind forcing. (C),(D) The same characteristics as in (A) and (B), respectively, but
after an additional 3 months of symmetric cyclonic wind forcing.
Figure 4. (A) The FC anomaly
(solid blue line) from observations and SSHG (red dashed line). (B) The FC anomaly
(solid blue line) from reconstruction and SSHG (red and yellow bars) as defined by P&J. The thick black line depicts the sea ice volume (km^3/year) anomalies from Hilmer and Lemke [2001]. Vertical axes show units of SSHG (x10^-6), sea ice volume anomalies (km^3/year), and the FC anomalies (km^3/year).
Figure 5. Sea ice concentration (SIC) in the GIN Sea averaged for the two CCRs: 1980-1983 and 1989-1997 (A),
and two ACCRs: 1984-1988 and 1998-2000 (B).
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