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UNIVERSITY OF GOTHENBURG
Department of Earth Sciences
Geovetarcentrum/Earth Science Centre
Recent evolution
of the Polar Mixed Layer,
the Cold Halocline
and Atlantic Layer Properties
of the Eurasian Basin
of the Arctic Ocean
Camilla Sten
ISSN 1400-3821
Mailing address
Geovetarcentrum
S 405 30 Göteborg
B632
Master of Science (Two Years) thesis
Göteborg 2011
Address
Geovetarcentrum
Guldhedsgatan 5A
Telephone
031-786 19 56
Telefax
031-786 19 86
Geovetarcentrum
Göteborg University
S-405 30 Göteborg
SWEDEN
ABSTRACT
Along track hydrographic data from the LOMROG II 2009 expedition from Svalbard to the North Pole
and Lomonosov ridge is compared with available post 1990’s datasets from the region in order to
present the evolution of the arctic halocline and mixed layer properties following the prominent
changes in the upper Arctic hydrography in the early 1990's. The 2009 data reveals an expansion of
the cold polar halocline now extending across the Amundsen Basin to the Nansen-Gakkel Ridge. Here,
the data document a large increase of the freshwater loading and potential energy implying an
increased geostrophic flow towards Greenland.
In addition, in 2009 the cold halocline layer above the Lomonosov ridge is weak with a strong
halocline on both sides of the ridge. It is during the recovery phase of the halocline in the early 2000
that this topography related shape occurs. Before the retreat it was absent.
Furthermore, a brief overview of diffusive processes is included that show evidence of salt fingering.
Finally, the results are evaluated with respect to changes in the large scale atmospheric circulation
described by the Arctic Oscillation and the North Atlantic Oscillation, also the Pacific Decadal
Oscillation has been taken into account.
Keywords: Arctic Ocean, cold halocline layer, Atlantic layer properties, warm pulses, large scale
atmospheric systems.
SAMMANFATTNING PÅ SVENSKA/SUMMARY IN SWEDISH
Data från LOMROG II expeditionen 2009, från Svalbard till Nordpolen och Lomonosovryggen, är
jämförda med tillgängliga data från 1990 till 2008 från samma region. Detta för att kartlägga
utvecklingen av den arktiska haloklinen och det välblandade lagrets egenskaper, samt undersöka om
dessa följer de temporala förändringarna i den övre delen av den arktiska vattenkolumnen. 2009 data
uppvisar en utbredning av den kalla polara haloklinen, vilken nu breder ut sig över Amundsen
bassängen helt ner till Nansen-Gakkelryggen. Här påvisar data en minskning av färskvatten och
potentiell energi, vilket medför ett baroklint geostrofiskt flöde mot Grönland.
Dessutom, år 2009 är den kalla haloklinen ovanför Lomonosovryggen svag med en stark och stabil
haloklin på båda sidor. Det är under återhämtningsfasen av haloklinen i början av 2000-talet som
denna topografi relaterade form uppkommer. Innan reträtten var fanns den inte
Vidare har en enkel översikt av de dubbeldiffusiva processerna och bevis för saltfingring inkluderats.
Till sist är resultaten utvärderade med hänsyn till det arktiska indexet, det nordatlantiska indexet
samt till motsvarande Stillahavsindex.
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PREFACE
During the last ten years the Arctic Ocean has become a major subject for the climate debate. Evidence
for warm water anomalies in the Atlantic Layer, and retreat of the Cold halocline Layer was published
as potential harbingers for rapid climate change. By adding new data sets that story can partly be
modified. Here, we present a set of investigations highlighting the temporal evolution of key diagnostic
properties of the upper Arctic Ocean and illustrate to what extent these are linked to earlier data.
This thesis is the author’s degree project for her Master of Science exam, with a major in Physical
Oceanography. The work was carried out in 2009 and 2010. In the fall of 2009 the data analysis took
place at the Danish Meteorological Institute in Copenhagen, and then further work with writing and
compilation were carried out during 2010 at the University of Gothenburg. In this thesis we have used
the LOMROG II data set, which were conducted onboard I/B Oden in the summer of 2009 as starting
point for analyses and visualization.
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TABLE OF CONTENTS
ABSTRACT ...................................................................................................................................................................... 1
SAMMANFATTNING PÅ SVENSKA/SUMMARY IN SWEDISH............................................................................ 1
PREFACE ......................................................................................................................................................................... 2
TABLE OF CONTENTS ................................................................................................................................................. 3
TABLE OF ACRONYMS AND SHORTENINGS ........................................................................................................ 4
1 INTRODUCTION AND APPROACH ........................................................................................................................ 5
2 AREA DESCRIPTION ................................................................................................................................................. 6
3 DATA SETS AND VISUALIZATION ........................................................................................................................ 7
4 RESULTS ....................................................................................................................................................................... 9
4.1 CLIMATOLOGICAL REFERENCE DATA, PHC 3.0 ........................................................................................................ 9
4.2 TRANSECT ONE, TIME DEVELOPMENT ..................................................................................................................... 10
4.3 TRANSECT TWO, TIME DEVELOPMENT .................................................................................................................... 14
4.4 THE LOMROGII DATASET; AUGUST, SEPTEMBER 2009 ........................................................................................ 16
4.5 TEMPERATURE AND SALINITY TIME SERIES ............................................................................................................ 17
4.6 ATLANTIC WATER VARIABILITY ............................................................................................................................. 18
4.7 THE COLD HALOCLINE LAYER................................................................................................................................ 20
4.7.1 Time series of the cold halocline layer ........................................................................................................... 22
4.7.2 Diffusive Staircase in the cold halocline layer ............................................................................................... 23
4.8 FRESH WATER CONTENT AND GEOSTROPHIC FLOW PATTERN .................................................................................. 26
5 CONNECTION TO THE LARGE SCALE ATMOSPHERIC SYSTEMS ............................................................ 29
6 DISCUSSION ............................................................................................................................................................... 31
7 CONCLUDING REMARKS ...................................................................................................................................... 31
8 ACKNOWLEDGMENTS ........................................................................................................................................... 32
9 REFERENCES ............................................................................................................................................................ 32
3
TABLE OF ACRONYMS AND SHORTENINGS
AB
Amundsen Basin
AO
Arctic Oscillation
AOA
Artic Ocean Atlas
AW
Atlantic Water
BSW
Barents Sea Water
CB
Canadian Basin
CHL
Cold Halocline Layer
CTD
Conductivity Temperature Depth meter
EWG
Environmental Working Group
FWC
Fresh Water Content
GPS
Global Positioning System
LHW
Lower Halocline Water
LOMR
Lomonosov Ridge
MB
Makarov Basin
NAO
North Atlantic Oscillation
NB
Nansen Basin
NGR
Nansen-Gakkel Ridge
NPEO
North Pole Environment Observatory
OHC
Ocean Heat Content
PDO
Pacific Decadal Oscillation
PHC
Polar science center Hydrographic Climatology
PML
Polar Mixed Layer
SCICEX
SCientific ICe EXpeditions
SIM
Seasonal Ice Melt
SSXCTD
Submarine eXpandable CTD
TU
Turner Angle
UHW
Upper Halocline Water
WOA
World Ocean Atlas
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1 INTRODUCTION AND APPROACH
During the late 90’s and early 2000’s the Arctic Ocean and its adjacent seas became a major subject for
the climate debate. Evidence for warm water anomalies in the Atlantic Layer, and a simultaneous
retreat of the Cold halocline Layer was published as potential evidence for rapid climate change
[Serreze et al. 2000, Morison et. al, 2000]. Nowadays, we can partly modify that story due to the
increased amount and frequency of available data making new analysis addressing different
timescales. The main purpose of this study is to evaluate the recent evolution of the upper ocean
hydrography in the Eurasian Basin of the Arctic Ocean and to identify key controlling factors possibly
linked to climate change.
With offset in the new LOMROG II 2009 hydrographic dataset, a systematic comparison of halocline
strength and Atlantic Layer properties with post 1991 cruise data is performed. This recent evolution
is referenced to the climatological hydrography of the Arctic Ocean using the PHC 3.0 climatological
reference data base (further explained below). The Focus is the uppermost and intermediate parts of
the water column including the warm and high saline Atlantic water, the cold halocline structure and
the polar mixed layer frontal system along the Lomonosov Ridge. Additional analyses are performed in
order to map the geostrophic flow pattern highlighting the close correspondence between, the
gradients and the ocean freshwater storage. The topographic steering for different layers in the
stratified Arctic Ocean is also discussed and results are evaluated with respect to changes in the large
scale atmospheric setting as described by the Arctic Oscillation (AO), the North Atlantic Oscillation
(NAO) and the Pacific Decadal Oscillation (PDO). Finally, a more novel investigation of the LOMROG II
dataset is presented including analysis of the cold halocline layer. This includes a brief summary of
double-diffusive unstable processes likely active in the water column.
To facilitate a consistent and robust comparison of the available post 1990 cruise data, a common
coordinate system has been created using simply the distance of individual stations to Longyearbyen,
Svalbard. Observations are found to cluster on two transects and are split accordingly and projected
onto two cross ridge transects using the calculated distance. Such section based analysis of available
data is novel.
This thesis spans a wide range of subjects using a variety of methods analysis. The purpose has not
been to integrate the results in an overarching conclusion. Instead discussions are integrated in the
individual result sections. The chapter of discussion summarizes the findings and offer a perspective
by linking some findings to known atmospheric variability, as reflected by the AO and NAO.
The thesis focus on questions concerning structure of and linkage between the geostrophic flow, the
Atlantic layer characteristics and the cold halocline layer in the study area of the Arctic Ocean.

How has the recent evolution of the Atlantic water properties developed from previously
reported extraordinary high temperatures in the inflow branches in the early 1990’s [Serreze
et al. 2000, Morison et. al, 2000]?

Do the new LOMROG II 2009 dataset support continued recovery of the halocline structure
from its retreat in the early 1990’s and later partial recovery [Steele and Boyd, 1998; Boyd et.
al., 2002; Björk et.al, 2002]?

What spatial and temporal structures can be found in the geostrophic current field based on
available hydrographic data? Are these expected, and how well does the LOMROG II dataset
correlate with the climatological reference?
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2 AREA DESCRIPTION
The Arctic Ocean is a rather enclosed sea, connected by Bering Strait to the Pacific Ocean and by Fram
Strait to the Atlantic Ocean, also connected in the East to the Barents Sea via Norwegian and Greenland
seas. The Canadian Archipelago allow near surface outflow in west through Baffin-Bay [Jones, 2001].
About two thirds of the entire area, 9.4 x 106 km2, is constituted of deep ocean, and one third shelf
areas [Aagard, Coachman, Carmack, 1981]. The Arctic Ocean water masses are comprised by the warm
and salty (S>34.8, T>0) North Atlantic water, entering through Fram Strait and Pacific water, less salty,
cold (S<33, T<-0.5) and rich in nutrients, enters the Arctic Ocean through the shallow (50m) Bering
Strait. Furthermore, the river inflow to the Arctic Ocean shelf areas is significant, and important for the
upper ocean stratification. The river inflow does mainly come from the large Siberian rivers and
Mackenzie, see figure 1 for location. The amount of freshwater in the upper ocean varies seasonally
due to summer ice melt and season variable river runoff [Aagard, Carmack, 1989].
Figure 1 Schematic circulation of surface water (grey arrows) and the Atlantic Layer plus Upper Polar Deep
Water to depths of about 1700 m (black arrows) in the Arctic Ocean. The straight arrows represent the mouths
of major rivers. Adapted from Peter E. Jones (2001).
The Arctic Ocean is divided into two major basins the Canadian Basin and the Eurasian Basin, by a
submarine ridge, the Lomonosov Ridge (LOMR) rising approximately from 4000m to 1700m below the
surface. It outcrops from Greenland, touches the North Pole and reaches all the way to the Siberian
continent. Furthermore the Eurasian side, which this study is concentrated to, is divided into two sub
basins; the Nansen Basin (NB) and the Amundsen Basin (AB). These are separated by the less
prominent Nansen-Gakkel ridge (NGR). The Canadian side is separated into two basins; Makarov Basin
(MB) and Canadian Basin (CB), divided by Alpha-Mendeleyev ridge, figure 1 and 2.
The Eurasian Basin of the Arctic Ocean is characterized by a surface Polar mixed layer (PML) reaching
about 50m, where a pycnocline, usually called the cold halocline layer, separates the surface waters
from the inflowing warm Atlantic water (AW) [Rudels, 2004]. The term halocline is here misleading,
because it indicates a transition zones between two well defined water masses, whereas this halocline
in reality is comprised by several water masses from different sources. Even though all are colder than
0oC and layered above each other with the less saline on top. This halocline is typically 150m thick
[Rudels, 2004]. The halocline can be divided into two parts. First the uppermost, thinner, almost
isothermal halocline where the salinity increases sharply [Aagard, Coachman, Carmack, 1981], Below
this the lower halocline layer is located; where the salinity is approximately constant whereas the
temperature increases with depth, from freezing to above zero degrees (Tf<T>0oC) [Rudels et al., 1996].
In this study the halocline is referred to as the depth interval between the PML and the AW.
The turbulent activity in the halocline is low, and the vertical upward heat flux is mainly due to double
diffusive processes. [Rudels et al., 1996]. Situated below this is the cooler and fresher Upper Polar Deep
Water (UPDW) ( 0o C    0.5o C,34.85  S  34.9 ), extending to about 1700m, typified by a negative   S
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relation. Beneath is deep water and bottom water located which for the Eurasian Basin presents
temperature and salinity values of 34.94 and -0.95oC, respectively. The corresponding values for the
Canadian side are 34.97 and -0.53oC. [Jones, 2001]
As mentioned above water flows into the Arctic Ocean through Bering Strait, from here it turns either
left to the Siberian shelf area, before it crosses the Arctic Ocean towards Fram strait, or it turns right
following the Canadian archipelago to the Fram strait. The general surface flow pattern from Bering
Strait to Fram Strait is called the Trans Polar drift, which also is carrying sea ice. Through Fram Strait
and from Barents and Kara seas through St Anna Trough low salinity surface water also flows into the
Arctic Ocean. These waters can be traced along the Siberian shelf coast to the Laptev Sea where they
converge into the Trans Polar drift, returning to the Atlantic Ocean via Fram Strait. The intermediate
water has a different flow pattern, which are divided by the basins causing an anti cyclonic current in
each basin, with an in and out flow to the Atlantic through the Fram Strait. [Jones, 2001] For illustration
see figure 1.
3 DATA SETS AND VISUALIZATION
In this study a variety of datasets with different origins are used to make a spatial and temporal
comparison to the new LOMROG II data, and to the PHC post 1990 climatological reference data. Data
is split on subjective basis in order to obtaining two distinct time series of sections, which spans the
area of interest. Large parts of the hydrographic datasets are collected during summer, in August and
September. Only one dataset from May and one from July are included. The dataset is based on
expeditions using surface ships, as well as submarines and data from portable equipment flown out to
the measuring site by helicopter or ski plane. Figure 2 shows the spatial distribution of all the datasets,
concentrated into two main basin transverse transects.
Figure 2 Map with all datasets available for the present study. The legend to the right shows name and year of
the presented datasets. The numbers indicates which transects are called transect 1 and 2, respectively.
Observations from 1991 collected during a Swedish expedition in August and September with I/B
Oden was one of the first cruises that took place with purpose to map and investigate the basic
characteristics of the Arctic Ocean. A variety of data was collected; one transect from Svalbard crossing
NB, AB to LOMR and into MB. [Anderson et al. 1994] From this set of observations we could extract two
suitable transects matching our area of interest. See figure 2 for location.
SCICEX, Scientific Ice Expeditions, was a five years scientific research program and collaboration with
the U.S Navy and a variety of American Universities. During eight cruises, 1993-1999, and one
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workshop in 2000 the US Navy made access to a nuclear submarine for civilian research work in the
Arctic Ocean. [Morison et al. 1998].From the data collected on the Eurasian side of the Arctic Ocean,
suitable datasets from 1995, 1997, 1998, 1999 and 2000 was found, for locating the transects see
figure 2. SCICEX data is collected from a submarine equipped with an expendable CTD (SSXCTD),
which is well suited for this program since the technology makes it possible to measure temperature
and salinity profiles while the ship is still submerged. The SSXCTD probe package is launched at an
upward angle from the submarine, while trailing a signal back to the ship. When it has reached the
depth of 15m the probe is dropped and the CTD descends to 800m. Temperature and conductivity are
measured and the depth is calculated from the elapsed time and raise- and fall rate. Since the CTD
probe is dropped when the instrument is situated at 15m depth, there are no surface measurements
[Morison et al. 1998]. This is inconvenient for some of our analysis, why submarine data have not been
considered in those cases. It is important to notice that surface data is missing, in the SICEX data
derived sections.
The available data from 2001 which is collected onboard I/B Oden during the Arctic Ocean expedition,
are spring or early summer data. For the present comparative study one fine gridded transect from
Longyearbyen northwards over LOMR could be constructed from these data.
The summer dataset of 2005 from the Healy-Oden Trans-Arctic Expedition (HOTRAX) was conducted as
a two-ship cruise with I/B Oden and USCGS Healy. Oden was used for Physical Oceanography
investigations, measuring temperature and salinity with a well proven Sea-Bird 911 CTD. [Björk et. al,
2007]. From this dataset one section could be included in transect 1. See figure 2 for location.
The NPEO data from 2008 was collected with ski plane which landed on the ice. A hole was drilled and
CTD measurements were carried out with a portable winch system, carrying a Seabird SBE-19 Seacat.
(The NPEO official website, 2010-01-08) Unfortunately most of these measurements are outside our
area of interest, even though it was possible to extract two useful fragment of our transects, shown in
figure 2.
The LOMROG II cruise started at Longyearbyen, Svalbard August 2009, and finished in September the
same year at the same location. The main purpose of the cruise was to collect seismic, bathymetric and
hydrographic data over and around the Lomonosov ridge, and especially over the intra basin.
Hydrographic data were also collected in two transect when heading northwards respectively
southwards to Longyearbyen [pers. comm Steffen Malskær Olsen,]. Here, using the hydrography focus is
set to these two transects and data from the intra basin has not been included. In figure 3 one can see
the two transects one east and the other west from the North Pole. The first transect used contains
station O02, O03, H05, H06, O04, O05 O06, H07, O07 and H10, henceforth called Transect 1. The other
transects is constructed of the stations H01, H02, H03, O16, H20, H18, H17, O15, H16 and H15,
henceforth called transect 2. Data was collected with two CTD:s, one ship mounted, and one portable
for helicopter use. This two CTD:s were calibrated to each other. The instrument precision for salinity
is about 0.002 PSU, whereas the calibration difference was smaller, i.e. negliable (same for
temperature). [pers. comm. Steffen Malskær Olsen].
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Figure 3 Oceanographic stations at LOMROG II expedition. Yellow (pink) marks indicate station carried out with
ship mounted equipment (portable helicopter equipment).
4 RESULTS
In order to obtain spatially comparable sections with a common axis all stations along each transect
are simply referred by the distance from Longyearbyen, Svalbard to the certain measuring sites, built
on the GPS-position of every station. A bottom topography plot, roughly along the middle of all
observations is inserted for easier location, and for visualization of the topographic steering.
4.1 CLIMATOLOGICAL REFERENCE DATA, PHC 3.0
As a climatologically and historical reference the PHC 3.0 database is used. This is a combination of
NODC's 1998 world climatology (WOA), the EWG Arctic Ocean Atlas (AOA), and selected Canadian
data provided by the Bedford Institute of Oceanography. [Steele, 2001]. The PHC is provided at 33
depth intervals, in monthly, seasonal and annual average; here only the summer average, covering the
period July to September is used, in order to have a comparable season.
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Figure 4 Temperature [oC] and Salinity [PSU] sections in the depth range 0-1000m respectively 0-300m for the
climatological PHC 3.0 along transect 1. The lowest panel shows the bottom topography along the transect. For
details of the x-axis see introduction. The vertical lines are the positions for the measuring stations. In both ends
of transect the data is extrapolated another 50km for easier reading of the plot, i.e. the first (last) station is
located at the first (last) dotted line.
From the gridded field two transects; running approximately along the 2 major transects in our other
consistent datasets was extracted, and horizontally spline-interpolated using the same routine as for
every dataset. In figure 2, the core temperature of the warm AW is about 1oC above the NGR and colder
over AB and MB. In the southern part of NB there is a small core warmer than 2oC. All over AB the
surface salinity is less than 32.5, and over NGR and NB less than 33. We will later see that this
reference situation is unusual compared to later sections, by having very high fresh water content.
Noteworthy is the absence of a salt front above the LOMR.
The second PHC transect (not shown) is very similar to the first transect. It is found a little colder
(<0.5oC) in NB and the AW tends to occupy a narrower (<50m) depth interval. In MB transect 1 is
about 0.5 unit fresher than in transect two.
4.2 TRANSECT ONE, TIME DEVELOPMENT
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Figure 5 Temperature for 0-1000m depth for all available datasets 1991 to 2008 in transect 1. The vertical
dotted line in each plot shows the location of the station. The bottom panel shows the bottom topography along
the transect.
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Figure 6 Salinity for 0-300m depth for all available datasets 1991 to 2008 in transect 1. The vertical dotted line
in each plot shows the location of the station. The bottom panel shows the bottom topography along the transect.
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The 1991, core temperature over NGR reaches 1.5oC and further north the 1oC water is spread all over
the area, see figure 5. Another core with 1.5oC warm water is located over the LOMR. This might well
be the remains of a former warm water anomaly, travelling around the Arctic Ocean with the currents
and now heading southwards again, towards Fram strait. In the salinity field (figure 6) a pronounced
fresh water front is located from MB, confronting the LOMR.
In 1995 the layers representing the core the Atlantic water are significantly warmer than 1991. Over
the NGR it reaches 1.5oC at 300m depth. The Amundsen basin is a little colder, but again a pool of
warmer water is located over the LOMR, as also seen in 1991 but shifted to the Eurasian side of the
ridge. Here the AW over the ridge occupies a larger depth interval reaching both deeper and closer to
the surface. Even though there is an absence of surface data, a salinity front (figure 6) is identified
above the LOMR. With the lowest salinity at about 33.3 PSU which is remarkably saltier than 1991
data in the same depth interval. Since the water at the base of the PML is saltier than 1991, it can be
assumed that also the surface water is saltier.
In 1998 the core of 2oC warm Atlantic water, over the NGR penetrated even deeper. Data from 1991
and 1995 showed evidence of an isolated warm anomaly in the Atlantic layer above the LOMR. This
anomaly is no longer pronounced, instead a near isothermal core spread from the Amundsen Basin
with a core of Atlantic water temperature of about 1.5oC. This water is indeed crossing the LOMR,
penetrating into the MB. The 1998 salinity section mirrors the conditions in 1995 (figure 6), but with a
tendency of ongoing salinification of the salinity front over the LOMR. The less salty Pacific water type
on the Canadian side is saltier and the front less pronounced, possibly due to a deepening of the
freshwater entrained layers over the LOMR.
The 1999 dataset is collected in May, and may be considered to present late winter conditions,
compared to the most of the other data sets which are typically summer datasets, collected in August
or September. Since the 1999 data seems to show similar characteristics as prior years and do not give
any indication of a large seasonal signal, it is fully taken into consideration. A core of warm Atlantic
Water is still present over the NGR, there is also left an isolated warm anomaly in AB, and above the
LOMR, figure 5. A general tendency of cooling, compared to 1998, is seen for the core temperature,
even though relatively high values of about 1oC is found along the entire transect. A marked cooling
and freshening takes place from 1995 to 1999 in the CB resulting in a stronger vertical stratification,
but limited to the Canadian side of the ridge.
The SCICEX data from 2000 shows an even more extreme evolution in both the Atlantic and upper
ocean layer. Here (figure 5) 1.5oC warm Atlantic water can be traced throughout AB over LOMR. The
core with a temperature of 2oC extends to the middle of NGR. A new core with water of one degree can
also be observed in the CB. The Salinity field (figure 6) reflects very fresh upper ocean conditions in
MB and above LOMR, where the 33.0 PSU isoline has moved as far as to the NGR. Freshwater has
penetrated further down in the water column compared with previous years. This increases the
density stratification and may decrease the vertical heat flow from deeper regions through the cold
halocline layer. In conclusion, 2000 conditions show a strong increase of fresh water in the halocline
and very high temperatures below in the Atlantic layer compared to conditions since 1991. Also
considering the PHC pre 1990 climatology, 2000 is an extreme year.
In 2001, the upper ocean salinity front in CB is forced into the Eurasian basin, even though it has,
compared to 2000, partly retreated towards the CB. Furthermore, the salinity in MB is even fresher
suggesting a positive feedback with the strong stratification established in 2000. The temperature
distribution in the Atlantic Layer (figure 5) is very similar to 2000 conditions with the exception that
the 2oC core water is slightly deeper.
2005 conditions in the Atlantic Layer are notably colder than seen during the last decade. The core
temperature (figure 5) of 1oC has retreated from the MP and LOMR to only occupy the AB which in
2001 both contained 1oC water. We have to go back to 1991 to find similar cold conditions. However,
since data is missing from NB the situation here cannot be analyzed why a 2.5oC core temperature as
seen in 2001 can still be present here. The halocline layer (figure 6) is significantly fresher in the
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Eurasian basin than previous years, including 1991. Only CB demonstrates little saltier conditions as
compared to 2001 but still with a very well developed halocline and strong stratification. The 33 PSU
isoline associated with a sharp front is now located as far away as over NGR. In 2001 it was found in
the northern part of AB.
In 2008, only three stations are available to construct transect 1. Here, a sharp salinity front (figure 6)
is present just over LOMR, but low salinity conditions extend across the section. Furthermore, water
with a core temperature of 1oC is absent over the LOMR reflecting that a continued cooling of the
Atlantic later takes place. Since these data are very limited and coarse it is not possible to include the
2008 transect 1 in all the later analysis made, but used when possible.
4.3 TRANSECT TWO, TIME DEVELOPMENT
Figure 7 Temperature for 0-1000m depth for all available datasets 1991 to 2008 in transect 2. The vertical
dotted line in each plot shows the location of the station. The bottom panel shows the bottom topography along
the transect.
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Figure 8 Salinity for 0-300m depth for all available datasets 1991 to 2008 in transect 2. The vertical dotted line
in each plot shows the location of the station. The bottom panelshows the bottom topography along the transect.
In 1991, transect 2 which is located closer to Greenland, (figure 7) shows colder conditions of the
Atlantic layer compared to transect 1 with the core water of 1.5oC only present on the Svalbard side of
NGR. Transect 2 extends closer to Svalbard where one station in the NB still show an AW core
temperature of 2oC. The generally colder conditions towards Greenland are consistent with gradual
cooling of the Atlantic Layer along its route through the Arctic Ocean under the cold, stratified upper
layer. The well confined core of 2oC water in the NB could be evidence for a new warm water anomaly
entering the Arctic Ocean, which is discussed further below. The salinity field (figure 8) demonstrates
a larger amount of fresh water in the surface layer, with a sharp front 31.5 to 32.5 PSU located at the
Eurasian side of LOMR. A pool of fresher water is also located in AB, with salinity as low as 32 PSU.
Atlantic layer core temperatures in 1997 reach 2oC in the NB, covering the distance to NGR. Compared
with the same transect 1991, the warm core is more wide spread covering the entire NB and now
reaching to NGR. Core temperature of 1oC covers the whole AB which compared to the 1991 has
warmed, and the Atlantic water reaches deeper into the water column. However, the core depth of the
Atlantic water is about the same as 1991. Since surface data is missing analysis of the actual position of
surface salinity fronts are difficult to make. There is evidence for a surface gradient in salinity just
south of LOMR and a second frontal structure may exist, entering MB from the Canadian side of the
ridge. In general, compared to 1991 the upper layers are saltier in the Eurasian basin where the
freshwater present in 1991 is nearly absent. On the Canadian side of LOMR, surface layers do instead
show evidence of freshening. The evolution from 1991 to 1997 is supported by the comparisons made
on transect 1 from 1991 to 1998.
Topographic steering of ocean currents are clearly mirrored in the temperature section of 1997, with
major changes in Atlantic Layer water mass composition over NGR and again over LOMR. Also the
salinity section reflects that major horizontal changes are connected to the topography, which is most
pronounced crossing the LOMR.
15
A few stations could be projected onto transect two in 2008 (figure 7 and 8) where Atlantic layer core
temperatures in the Eurasian basin have cooled significantly from 1997 and only just exceed 1 oC. The
absence of warmer water towards NGR is striking and may reflect a colder or less strong inflow. The
salinity section shows resemblance with the 2005 conditions on transect 1, with the 33.5 isohaline
outcropping near the northern end of NGR.
4.4 THE LOMROGII DATASET; AUGUST, SEPTEMBER 2009
Figure 9 Temperature [oC] and Salinity [PSU] for Aug-09 (LOMROG II) data, transect 1. The vertical dotted line in
each plot shows the location of the station. The bottom panel shows the bottom topography along the transect.
The temperature and salinity of transect 1 is shown in figure 9. The core temperature is just above 1oC
all over Amundsen basin, ending with a large interleaving structure just on the Canadian side of LOMR.
The PML temperatures are near freezing and have very low salinity. The PML of the Canadian basin is
fresher still than the Eurasian basin and there is a front in salinity just above the ridge. Note also the
tendency to rose and compressed isohalines (isopycnals) in the Atlantic Layer above LOMR.
On transect 2, figure 10, the Atlantic layer core temperature of 1oC does not reach all the way to the
LOMR as seen on transect 1. It is obvious that the water cools during its circulation around the Arctic
Ocean and the warmest waters is to be found in the southern parts close to the inflow through Fram
strait. The warm water lens in the interior of AB may be water that has circulated around the basin
and is now heading back towards Fram strait again. Evidence for this is shown later (section 4.8 and
5). A sharp surface salinity front is located over NGR, which corresponds well with the locations of a
colder core temperature of 0.5oC in the Atlantic layer. The salinity front over the NGR is more
pronounced than seen in any of the datasets reaching back to 1991 on transect 1 and 2.
16
Figure 10 Temperature [oC] and Salinity [PSU] profiles for Aug-09 (LOMROG II) data, transect 2, otherwise same
as figure 9.
4.5 TEMPERATURE AND SALINITY TIME SERIES
A mean value for each basin and for the region of the LOMR has been constructed for every year, in
order to reveal travel times and pathways of anomalies (figure 11).
Figure 11 Temperature time series of mean values for each basin for the uppermost 1000m of the water column.
Measuring lack is shown by white areas.
The tendency for the southern basins to be warmer than the northern areas is apparent (figure 11).
Carmack (1995) report a pulse of warm water entering the Arctic Ocean around 1993. Traces of this
pulse, circulating the Arctic Ocean can be found by the millennium shift in AB and above LOMR. A weak
signal is also to be seen in MB, by the same time. The timescale for a warm pulse entering AB through
Fram strait, circuiting around the Arctic Ocean periphery until it reaches MB, is about seven years
17
[Kikuchi, 2005]. The results also consists well with the results of Björk, et al. (2010), Morison et al.
(2006), Kikuchi (2005) and Polyakov (2004) where the same conclusions regarding the warm pulses
was drawn. The inflow of the early nineties was not only warmer but, as suggested by Shauer et al.
(2002), also presented an enhanced volume. The pulse signal fits well with the AW circulation patterns
of Shauer et al. (2002). The subsurface AW circulation is strongly affected by the bathymetry, flowing
cyclonically along the basin margins [Polyakov, 2004]. An inflow of a new anomaly can be traced in
Fram strait in 1999, and further recorded in Laptev sea in 2004 [Dmitrenko et. al, 2009]. This
corroborates with the warm water entering NB in 1998-1999 as suggested in this study (figure 11);
which then might be found above LOMR in 2009 data. Besides this the 2009 data shows evidence for
retreat of the warm pulse from NB and AB where temperatures are now comparable to pre 1995
conditions.
Figure 12 Salinity time series of mean values for each basin, for the uppermost 300m of the water column.
Measuring lack is shown by white areas.
Because of the lack of surface data from many years it is more difficult to identify tendencies in the
salinity field. However, NB and AB contain in general less freshwater, and the low saline surface layer
is in general shallower (figure 12). A noteworthy trend is the increasing amount of freshwater in MB,
after 2001, also seen over the LOMR and AB, but here somewhat weaker during the last part of the
period. This result corresponds well with McPhee et al. (2009), who evaluate the amount of fresh water
on the Canadian side, and found a large increase of freshwater there. The increased freshwater loading
after 2001 can also be seen in NB where the layer of relatively low saline water is deeper, even though
this water still has higher salinity compared to the other basins.
4.6 ATLANTIC WATER VARIABILITY
The warm and salty intermediate AW plays a special role in the thermal balance of the Arctic Ocean.
Isolated from drifting ice and the atmosphere by a surface layer, the AW carries significant quantities
of heat. [Polyakov et al., 2005]. The variability of AW characteristics is studied in further detail based
on the heat content (The Atlantic Water mass is defined as the water column with a temperature
>0oC). This is done by computing the ocean heat content (OHC), which is defined by equation 1. The
level of integration is set to the interval where T>0oC, i.e. the AW.
18
∫
(
( )
( )
( ))
(1)
Here Cp is the specific heat capacity [J Kg-1 K-1], T is the Temperature and ρ the density.
As indicated by the temperature OHC decrease northward towards the Canadian basin in all presented
datasets, when the relatively concentrated heat in the inflowing core is spread, mixed and cooled in
the interior basins. All datasets in transect 1 (figure 13A) demonstrates a steep gradient over NB and
NGR, whereas the OHC is approximately constant over AB and decreasing further towards MB. At all
individual stations the OHC is larger than the climatology reference at the Eurasian side, even though
1991 is close to PHC values. Except for the 1991 data the temporal spreading was very small, with data
following close to each other over the Eurasian basin. Especially over NGR the variance is small
whereas the spreading is larger in AB, increasing over LOMR towards the CB. Here, the first datasets
shows decreasing OHC, whereas the more recent data shows an increasing trend. Although the time
variability is small there is a trend with increasing OHC in AB, and over NGR until 2001. When the first
anomaly retreated the OHC decreased again towards climatology. Topographic effects are however,
more pronounced in the analyzed data compared to climatology. An exception is 2005 which seemed
to have a low OHC at some sites and very high in others.
Figure 13 and Ocean Heat Content [MJ/m2] in the Atlantic Layer, i.e. integrated over the water column where
T>0 oC, for all available datasets, transect 1 (A), and transect 2 (B). The bottom topography along the transects is
shown at the bottom of each panel.
The 1991 transect 2 (figure 13B) gives further evidence for generally warmer conditions in the AB
after 1991 and onward compared to the climatology. This is consistent with the anomalies observed
by Carmack (1995), Shauer et al. (2002) and Dmitrenko et al. (2009) entering Fram strait during the
early 1990:s and further on during the 2000:s. While 2009 demonstrates a curve significant PHC,
although the OHC has increased slightly.
19
A complement to OHC the depth of maximum subsurface temperature is shown in appendix 1. There it
can be seen that the core temperature depth of the Atlantic layer always is shallower in the
climatological data.
4.7 THE COLD HALOCLINE LAYER
As a transition zone between the Polar Mixed Layer (PML) and the penetrating Atlantic water is the
halocline layer located. The cold halocline prevents upward heat flux from the warm Atlantic water
and thus helps maintain the sea ice cover. The term halocline can somewhat be misleading, because it
normally indicates a transition zones between two well defined water masses, whereas the Arctic
halocline in reality is comprised by several water masses from different sources. [Rudels, 2004].
As suggested by Rudels et al. (1996), the Cold Halocline Layer (CHL) is formed by mixed layer
convection down to ~120m north of Svalbard. This convective layer is here covered by advective (ice
melt and river run off) low salinity and cold (T<0.5oC) shelf water, forming a cold halocline structure
[Kikuchi et. al, 2004]. In the upper part of the Halocline (UHW) the salinity increases a lot with depth
while the temperature is nearly constant. Below this the lower halocline water (LHW) is located;
where the salinity increases from 33 to 34.2 PSU [Steele and Boyd, 1998] whereas the temperature
increases with depth, from freezing to 0oC (Tf<T>0oC) [Rudels et al., 1996], see figure 14.
Figure 14 Schematic picture of the Convective-advective CHL formation. Upstream (downstream) conditions are
shown as blue (black) lines. Tf = Freezing Temperature. For further explanation see text and Table of Acronyms
and Shortenings.
Evolution of the CHL in the Arctic Ocean has been an important topic among Arctic oceanographers
over the past few decades, starting with the pioneering work of Aagaard and Carmack (1981). It is still
a reasonably difficult area of investigation, and many theories on the formation and effects of the
halocline layers are suggested. This is partly because the halocline is not a specific water mass with a
stringent definition, but a transition zone of convection and mixing of different advective water masses
[Steele and Boyd, 1998]. By the same reasons many different definitions has been used, depending on
the coverage of the available data, which kind of data is obtainable and which analysis to be made.
According to Boyd [pers. comm.], and used in a variety of papers, the LHW can be defined as the depth
range with
PSU. This straightforward definition is convenient and applicable to the data
used in this study.
To assign the UHW the definition following Boyd et al. (2002) a salinity average for the depth range 7090m, is used. The fresh water signal from seasonal ice melt (SIM) typically extends no deeper than
65m, therefore a salinity average of 80m±10m makes an estimation of the salinity at the uppermost
part of the CHL. Low salinity in this range represents a stronger stratification, due to the increased
buoyancy difference between the AW and below where the salinity is relatively constant and always
greater than 34.9 PSU [pers. comm. Boyd].
Since the CHL is acting as a barrier to the upward mixing of heat it insulates the surface water from the
heat contained in the AW, this have profound effects on the surface energy and mass balance of sea ice
20
in this region. With an absent, or very weak halocline layer a significant reduction of the sea ice growth
during winter can be expected as a result of the upward mixing of heat from the warm Atlantic Layer.
[Björk et. al, 2002] With the above shown two definitions of the CHL, we should manage to visualize
the temporal and spatial variability of the halocline’s strength and existence.
The depth levels of the reference PHC database are 50m, 75m and 100m, which match poorly with the
depth interval for the calculations. Values have therefore been interpolated and care should be
exercised in the interpretations of these results, because the construction of the gridded PHC
climatology already includes averaging and interpolation.
Figure 15 Mean salinity [PSU] at 80m±10m for all available data in transect 1 (A) and transect 2 (B). The light
blue interval indicates the salinity 34<S<34.4, if the data is attaining this level the winter mixed layer shows
evidence of forming LHW, and a CHL is absent. The black horizontal line indicates S=34.8, which is the salinity
below the halocline. The bottom topography along the transects is shown at the bottom of each panel.
In transect 1 (figure 15A) the halocline appears weaker than the PHC reference at all locations, but in
general both climatology and observations show a strengthening towards LOMR and into CB. The
gradient is less pronounced in 1991 than in 2009. The stratification is in general notably stronger in
2009 and 2005, compared to the previous datasets. Furthermore, between 1991 and 1995 the
halocline was retreating from central AB to a weak or absent CHL all over the Eurasian side, this
tendency continued during the late 90’s and early 20’s. 2001 CHL has recovered back again in AB, this
recovery progressed to 2005. In 2009 CHL has reestablished coverage and strength comparable to the
conditions of 1991, except the narrow area over LOMR at the Eurasian Basin side. In addition, the
slope towards CB is increasing over the entire period; the stratification is consequently stronger than
at the Eurasian side. Close to LOMR, on the Eurasian side CHL is weaker, compared to more far away
both sides of the ridge. Nevertheless, the horizontal gradient associated with LOMR is larger in 2005
21
and 2009 than in the previous datasets indicating that the topographic steering has a larger impact on
the water column.
Transect 2 (figure 15B) demonstrates a slope in 2009 which behaves exceptionally similar to the
climatology, except for the increased salinity on the Eurasian side of LOMR. This topographic effect is
noticeable in the reference climatology data as well, even though it is remarkable more pronounced
2009. Likewise, the slope has increased with time; a comparison of 1997 and 2009 confirm significant
temporal changes.
Further on, in 2009 transect 1 CHL above LOMR is diminutive, accompanied a strong halocline on both
sides of the ridge, whereas this tendency is absent in 1991. During the recovery phase of the halocline
in the early 2000 this topography related shape occurs. Before the retreat, shown as 1991 and in PHC
reference, it was absent. Furthermore, the first southernmost stations 2005 (fig. 15A) show a very
weak or negligible halocline as discussed above, whereas the fifth and sixth station demonstrates a
very strong halocline. Here the water mass is very similar to the waters on the Canadian side in 2005,
why this might not be a clear case of the presences of CHL, but instead an event of Pacific water
intruding the Eurasian basin. The same signature is seen 2008 and 2009 in transect 2 (figure 15B).
Figure 16 Potential temperature (θ) [oC] and Salinity [PSU] data for Sept-91 (I/B Oden) and Aug-09 (LOMROG
II), from transect 1, shown in θ/S-space. A mean Salinity 70-90m and temperature in the same depth interval is
used. Used reference pressure is 80 dbars. The dashed line indicates the freezing point of seawater.
Moreover, a T/S plot (figure 16) yields further evidence of the state of CHL before the retreat was
convectively formed while it after the recovery indicates an advective formation. The 1991 data is
scattered in the expected area of CHL, except one data point to be found in typical AW (cold and salty),
located above NGR, and one data point showing typical Pacific water (warm and fresh) characteristics,
situated in MB. T/S curve for 2009 behaves significantly different; six data points demonstrate obvious
Pacific characteristics, reaching from MB well into AB.
Additional support documenting this difference is found by plotting the mean temperature in the same
depth range (see appendix 2). The 2009 data demonstrates a spatial variation contrary to 1991 data.
1991 the mean temperature over NGR and in AB was significant higher than the PHC, decreasing
towards the LOMR and the Canadian Basin.
4.7.1 TIME SERIES OF THE COLD HALOCLINE LAYER
With the same definitions of CHL a time series has been conducted in figure 17. An average salinity for
each basin has been computed in the mentioned depth range, and compared with each other.
22
Figure 17 shows time series for the mean Salinity at 70-90mts, for every basin. The light shaded horizontal
interval indicates the salinity 34<S<34.4, if the data is attaining this level the winter mixed layer shows evidence
of forming LHW, and CHL is absent. The dashed black horizontal line indicates S=34.8, which is the salinity below
the halocline.
This perspective shows further evidence of the timing of the retreat and recovery discussed above.
The results correspond well with the results of Boyd et al. (2002), Steele and Boyd (1998), and Björk et
al. (2002). They show that during the 90's the CHL retreated from middle of Amundsen basin over
LOMR and, in 1998 the CHL is absent all over the Eurasian basin. Moreover, in correspondence with
Björk et al. (2002), in early 2000's a recovery of the halocline is progressing back into AB.
Note also, in figure 17, that a well-developed CHL is always present in MB. Here the salinity decreased
during the 90’s, interrupted by a sudden increase in 2001. Decreasing again to the end of the record,
with the overall lowest values in 2009. Excluding the situation in NB the general situation of 2009 is
CHL of the same strength or stronger than in 1991 Furthermore, a trend towards stronger CHL exists
2001 to 2009.
4.7.2 DIFFUSIVE STAIRCASE IN THE COLD HALOCLINE LAYER
Double diffusion is defined by convective motions caused by different rates of diffusion of heat and salt
in the ocean. In the Arctic Ocean cold and fresh water masses are located above warm and salty water
in the Atlantic water. This causes unstable temperature stratification in the halocline. The molecular
diffusion of salt is about hundred times smaller than for heat, therefore the AW, will lose heat upwards
faster then it loses salt. The water above the interface, i.e. halocline water becomes lighter while it
gains heat and tends to rise, whereas the water below the interface gets denser and tends to sink. The
result is small homogeneous convective cells, separated by thin but strong gradients, which appears
like a stair, (thermohaline staircase) [Shi, 2006]. Another double diffusive situation is when warm and
salty water masses are located above cold and fresh water. The water above the interface becomes
colder and sinks, and the water below the interface becomes warmer and rise. The result is small scale
(cm) finger shaped water columns; Salt fingering [Shi, 2006]. In the Arctic Ocean where cold and fresh
water lies above the warm and salty Atlantic water, we generally expect to find evidence of the
diffusive regime.
To quantify the strength of the double diffusive processes, the stability ratio R which is the
temperature- and salinity gradient ratio normalized by the salt contraction coefficient β, [PSU-1] and
the heat expansion, α [oC-1] is taken into account (eq. 2) [Rudels et al. (1999), Bianchi et. al (2002),
Thorpe (2007)].
(2)
23
Temperature and salinity have opposite effects on the density. The tendency to have active double
diffusion is as largest when R  1 [Thorpe, 2007] i.e. when the columns are weakly stable. The Turner
angle (TU) defined in eqution 3, can be used as an indicator of double diffusive activity. The TU is
(
) , [You, 2002] although it has many advantages over using the
related to the Rρ by
stability ratio. For example may two observing points have the same R  , even though the stratification
in one case is stable and in the other unstable, and they will have different TU [Bianchi, 2002].
(
) {
(3)
-90⁰< TU< -45⁰
Diffusive regime, T is unstable. (0 < Rρ <1)
-45⁰ < TU < 45⁰
Temperature and Salinity have the same diffusion rate, why the water column is
doubly stable.
45⁰ < TU <90⁰
Salt finger favorable, S is unstable. (1 < Rρ < ∞)
All other values represent an ordinary statically unstable stratification [You, 2002]. Figure 18
illustrates the definition of the TA.
Figure 18 Illustration of the definition of the Turner Angle, with corresponding R indicated, following You
(2002).
Thus, the TU unambiguously express the exact stability at a certain depth. TU together with salinity
and σ-profiles give sufficient evidence for double diffusive processes in the LOMROG II data.
Clear evidence of a diffusive staircase can be seen in most of the profiles. The stairs become more
prominent towards north, as exemplified by station H15 (see figure 19), from the Canadian side of the
LOMR. In general 2009 profiles are much smoother compared to observations by Rudels et al. (1999)
who used data from 1991 and 1996, when there was much more developed double diffusive layering
especially in AB. Due to Rudels et. al (1996) in a case of a weak or absent CHL, the temperature steps
would be sharper. Larger double diffusive fluxes are to be expected, whereas a strong halocline will
keep the heat. This implies the smoother profiles in 2009, since the halocline in AB and above LOMR is
stronger.
24
Figure 19 Salinity and Temperature profiles for station H15, in LOMROG II data.
TA has been plotted with respect to the depth, figure 20 left, to give further evidence for diffusive
processes are active in the depth range 170-270m, i.e. the upper part of the CHL. As additional
information the buoyancy frequency (N2) has been plotted in the same depth interval. It is clear that
within a saltfinger interval the stability is close to zero and increase markedly in the diffusive
intervals.
25
Figure 20 Left: Turner Angle. Light blue (light green) indicates saltfinger (double diffusive) regime. Black lines is
the temperature and salinity profiles for the station and depths equally, for scaling see figure 20. Right: Buoyancy
Frequency (N2). Station H15 in Lomrog II dataset.
4.8 FRESH WATER CONTENT AND GEOSTROPHIC FLOW PATTERN
The freshwater content (FWC) of the water column is defined as
∫(
)
(
)
(4)
where Sref equals 34.8. The integration interval is from the depth of the salinity reference to the surface
[McPhee et al., 2009]. As shown above surface data from several years and locations is missing. To be
consistent, all stations lacking surface data have therefore been excluded from this analysis. Using
stations with missing surface values would cause a large under estimation of FWC since the uppermost
meters contain a large quantity of freshwater.
The Geostrophic flow field is computed using the classical approach, starting from the baroclinic
geostrophic balance (equation 6)
(5)
Considering the pressure as hydrostatic, deriving with respect to Z, the result is the equation for
thermal wind (equation 7)
(6)
Integrating from the sea surface
to a level of no motion below the halocline, Z=300m, the
geostrophic flow pattern has been computed.
Following Björk, et al (2001) a potential baroclinic flow field can be estimated using the vertical
integrated potential energy, P, relative a reference water mass
(
) Where. Sref =34.8,
26
and Tref is an approximated temperature Tref  T
Z ( Sref )
 0.6o C ,
assuming v=0 at ρ=ρref is set close to
the deep water properties.
∫
(7)
The baroclinic flow, Qρ is directly related to the Potential Energy through the relation
where f is the Coriolis parameter. The potential baroclinic flow field should be a rough estimate of the
baroclinic transport that is superimposed on an unknown barotropic flow field. For the computations
fixed integration interval,
m has been used. The advantage of this fixed intergration depth
is that the calculations become less sensitive to small temporal and spatial variations in the deep water
which can corrupt the computations [Björk et al., 2001].
Earlier studies have estimated the FWC to be in the range 0m and 10m, increasing towards CB. The
FWC has increased rapidly in CB, and studies shows that the values here are much larger than the
climatological reference [McPhee et al., 2009]. This signal dominates also the results of the present
analysis (figure 21).
Figure 21 Fresh Water Contents (FWC), shown in meters of the water column, considering transect 1(A) and
transect 2 (B). FWC was calculated only for data sets where surface data was available. The bottom topography
along the transects is shown at the bottom of each panel.
At all locations climatology shows a larger FWC than data from Transect 1 suggests, figure 21A. 1991
data showed a comparable small slope, whereas the slope increased remarkable into the Canadian
basin. The amount of fresh water between the two ridges has increased with time more than above
NGR or LOMR. It is also shown that the amount of fresh water has increased much in MB over time.
2009 data shows a pronounced bending over LOMR. This decreasing bending was presented in 2005
as well, although the rising in MB was less pronounced.
In the second transect (figure 21B) the 1991 data contained more fresh water than PHC in NB, this has
during the years retreated to a level below PHC in 2009. The sudden increase over NGR in 1991 is
probably due to the data projection to a common distance axis, otherwise the slope in this data set is
27
minor than the climatology and 2009 data equally. In addition, 2009 data demonstrates a behavior
similar to climatology. Even though the shape of the curve is more pronounced, particularly in the
southern parts of AB. The northern parts of AB and above LOMR show values close to climatology.
Moreover, the shape of the curve is displaced towards LOMR, as the most over NGR and in the
southern parts of AB where the difference between 2009 data and PHC is large.
On the Canadian side of LOMR there is a distinct increase over time. This fits well to the results of
McPhee et al. (2009), who found that the amount of freshwater in the Canadian Basin had increased
dramatically in 2008 compared to the Climatology (PHC).
Figure 22 Geostrophic current, Transect 1, from the surface to the level of no motion, set to 300m. Positive, red
(negative, blue) currents are directed towards Greenland (Russia). The bottom topography along the transect is
shown at the bottom of each panel.
Considering the Geostrophic currents, the climatological values was close to uniform, with magnitudes
<5cm/s over NB and NGR (not shown). Comparing AB and above LOMR the flow pattern proved more
inconsequent pattern, with some irregularities, this is suggested to be dependent on the influence of
the ridge. In all datasets the current changes direction over LOMR, which were not the situation of the
PHC reference. The 2009 data were showing a displacement of the irregular peaks compared with
previous data, towards Svalbard, this tendency might be connected with the increased gradient of FWC
over the middle and north parts of AB (compare figure 21A). In general a large fresh water loading and
FWC gradient implies large baroclinic geostrophic currents. Whereas in areas where the current
change direction or is diminutive theFWC is small, and the rise is low. The general pattern of figure 22
suggests that the baroclinic current is influencing the water column to approximately 170m, but it can
be traced to about 250m.
28
Figure 23 Potential Energy, measured in [m3/s2] for transect 1. The bottom topography along the transects is
shown at the bottom of the panel.
At all occasions the gradient of potential energy is decreasing or flattens out immediately over the
LOMR (figure 23). The bathymetric change caused by the ridge is affecting the whole water column, so
even the surface water. In all datasets the potential energy increases towards north, and is in general
higher on the Canadian side than on the Eurasian. The curves of potential energy is shapely similar to
the FWC curves.
5 CONNECTION TO THE LARGE SCALE ATMOSPHERIC SYSTEMS
The warm isolated anomaly found in the AB in the late 90’s is probably the warm pulse entering the
Arctic Ocean in 1993, as suggested by Björk et al. (2010), Morison et al. (2006) and Kikuchi (2005).
Kikuchi (2005) suggest a timescale of about seven years for circulation the Eurasian basin, which is
suitable for the observed anomaly in AB. The flow pattern of these pulses as shown by Shauer (2002),
and corresponds to the geostrophic flow pattern in the Eurasian basin, also mentioned by Jones (2001),
and shown in this work in figure 1.
In 1998 a new inflow of warm Atlantic water entered the NB through Fram strait as indicated in figure
11. As suggested by 2009 data there might be a remnant signal over LOMR, using the same timescale
of seven years [Kikuchi, 2005]. This pulse is also documented by Polyakov et. al (2005) who found the
warm anomaly at the border to the Laptev sea shelf in 2004. Besides this, 2009 data shows evidence
for retreat of the previous warm pulse and temperatures are comparable to pre 1995 conditions.
29
Figure 24 shows North Atlantic Oscillation (NAO) (upper), Arctic Oscillation (AO) (middle) and Pacific Decadal
Oscillation (PDO) (lower) between January 1990, and January 2010. The red line is the four month running mean
value. The dashed vertical lines indicate January 1993, February 1997, and December 2006.
Also shown by Polyakov et al. (2004) it seems like these warm pulses are connected to the large scale
atmospheric circulations, such as the Arctic oscillation (AO) and the North Atlantic oscillation (NAO)
(figure 24). They suggest that the AW variability is dominated by multidecadal fluctuations in a larger
perspective with a time scale of 50-80 years. Here, even the Pacific decadal oscillation (PDO) is taken
into account. A high positive phase of the NAO is associated with stronger winds from the south over
Barents Sea and Fram strait. This influences the mean surface wind field and therefore the mean sea
level pressure. Whereas, a negative index is characterized by a local maximum in the mean sea level
pressure which is leading to light and variable winds. Evidently there is an anomasly high AO index in
January 1993 and in February 1997. It is also found that this anomaly is surrounded by phases with
negative indices. Comparing with the other indices this occurs in a situation during a positive phase of
NAO and when PDO has values close to zero. Perhaps this particular situation is trigging and allowing
warm pulses to entering the Arctic Ocean by Fram strait.
Another prominent short deviant positive phase in the AO is seen in December 2006, which also
corresponds with a positive phase of the NAO, and a PDO close to zero. Consequently, a new warm
pulse originating from this anomaly might be located in AB and over the LOMR in approx 2014. We
cannot see this pulse entering NB in 2007. This is simply due to lack of data from this particular year,
and the pulse is lost in the interpolation routine. Despite this, Shauer (2008) shows a warmer core
temperature of the inflowing AW through Fram strait in 2006. If this pattern is true and traceable than
these intrusions of warm Atlantic water is likely a natural phenomenon due to natural modes of
inertial variability. Further studies require investigations on why these spikes in the AO-index are
occurring, and what cause them.
The warm pulse that has circulated around the Arctic since 1993 was warmer and obtained a large
volume then inflowing Atlantic water other years (Shauer 2002). This influence is obtaining a larger
part of the water column and may affect the halocline in AB, compare figure 17, where evidence for a
30
nonexistent CHL in AB is shown the current years. The AW is saltier than UHW, which means that the
possibility for the winter mixed layer to form LHW is larger when more Atlantic water is present,
which results in an absent CHL. If a larger amount of heat is present in the water column, the heat flow
through the halocline is potentially larger. This combined with the weakening of the CHL may affect
the ice cover and triggering melting processes. If the above explained effect of atmospheric circulation
systems on the inflowing AW is true, we can apply this on the CHL and connect the strength of it to the
characteristics of AW inflow.
6 DISCUSSION
Regarding the results of the CHL (figure 15), the general discussion is concentrated on whether it is
convectively- or advectively formed. Considering figure 16, in 2009 an event of Pacific water might be
present at the Atlantic side. What this intrusion is caused by we cannot explain here, but it is clear that
2009 conditions in many aspects are more topographically steered in the area of the ridge. This new
pattern can be a consequence of the possible intrusion of Pacific water. A certain conclusion cannot be
drawn here, but we suggest further, more profound, studies were nutrient samples is taken into
account. This might yield further insight to this event.
McPhee et al. (2009) shows that the amount of freshwater in CB rapidly increases, with high anomalies
in FWC, dynamical heights and also geostrophic flow and transport. This pattern is caught in our
analysis as well, figure 21. Close to the border to the Canadian side FWC has increased markedly.
LOMROG data exhibits larger variability than climatological data, concerning potential energy and
FWC. The topographical steering is more dominant here than in PHC reference system. Also it can be
noted that the topographical steering is increasing with time. The increase of fresh water between
2001 and 2005 is well pronounced. Comparing this to the intrusion of the isolated warm Atlantic
anomaly and the AO index, the timescale corresponds with the time when the pulse is expected to
reach LOMR, and the northern parts of AB. May this raise of potential energy and FWC be a cause of
the circulating pulses? Since the available surface data is limited a firm conclusion is hard to make,
even though it seems like the results are directed towards this hypothesis. The baroclinic flow pattern
showed in figure 22 corresponds to the expected, both taking FWC and potential energy into account,
but also referring to Jones (2001) and Shauer (2002). It can be concluded that the topographical
steering of the baroclinic geostrophic flow has increased with time.
Comparing the baroclinic geostrophical flow field with the PHC reference is hard. The PHC shows in
general a much smaller magnitude at all occasions, and the topographical steering is diminutive. The
PHC database is good to use as a reference but since the data has been both horizontally and vertically
interpolated due to the desirable grid field, small changes due to for example bathymetry are lost and
the reference become misleading in these cases.
7 CONCLUDING REMARKS

In 2009 the cold halocline layer above the LOMR is weak with a strong and stable halocline on
both sides of the ridge. It is during the recovery phase of the halocline in the early 2000 that
this topography related shape occurs. Before the retreat it was absent.

LOMROG II data exhibits a larger variability than the climatology, concerning potential energy
and FWC. The topography steering is more dominant here than in PHC reference system. Also
it can be denoted that the topographical steering is increasing with time. The increase of fresh
water between 2001 and 2005 is well pronounced. Comparing this to the intrusion of the
isolated warm Atlantic anomaly and AO index, the timescale corresponds with the time when
the pulse is expected to reach LOMR, and the northern parts of AB.

Considering the double diffusive heat flow the 2009 profiles are in general much smoother
compared to observations by Rudels et al. (1999) who used data from 1991 and 1996, when
there was much more developed double diffusive layering especially in AB. Due to Rudels et. al
31
(1996) in a case of a weak or absent CHL, the temperature steps would be sharper and larger
double diffusive fluxes are to be expected, whereas a strong halocline will keep the heat. This
implies the smoother profiles in 2009, since the halocline in AB and above LOMR is stronger.

Our results on the warm pulses entering the Arctic Ocean corroborates with previous
publications. It seems like these warm pulses are connected to the large scale atmospheric
circulations, such as the Arctic oscillation and the North Atlantic oscillation.
8 ACKNOWLEDGMENTS
The author wish to express her great thanks to the supervisors Steffen Malskær Olsen at the Danish
Meteorological Insitute (DMI) and Göran Björk at the University of Gothenburg (GU) for support, new
perspectives and help. A great thanks is also directed towards the Danish Continental Shelf Project, to
Christian Marcussen, researchers and crew onboard I/B Oden during the LOMROG II cruise to the
Lomonosov Ridge off Greenland. Special thanks are directed to Leif Toudal Pedersen (DMI), and to
Timothy Boyd for personal comments and help, also to Antonio Cuevas (UiB) for personal help.
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33
Appendix I
Depth of maximum subsurface temperature
Figure Depth of maximum subsurface temperature, 0<Z<1000m, for all available datasets, transect 1 (upper)
and 2 (lower). Shows even the bottom topography along the transect.
34
Appendix II
Mean temperature of the Depth range 70-90m
As additional information and for giving another perspective to subsistence of the CHL the mean
temperature in the depth range 70-90m has been plotted. It was distinguishing that data
demonstrated the contrary situation when comparing 1991 and 2009. 1991 the mean temperature
over NGR and in AB was significant higher than the PHC, decreasing towards the LOMR. What is
observed in more recent data is the opposite situation; a mean temperature well lower the PHC in AB,
increasing towards LOMR. In both cases the temperature is increasing on the Canadian side. The
intermediate datasets were also following this pattern.
Figure Mean temperature [oC] in the depth range 80  10 m for all available data sets, transect 1 (upper) and 2
(lower). The lowest plot shows the bottom topography along the transect.
Nevertheless transect 2 which show more southern conditions suggests the same tendency as transect
1 1991 exhibits a cold situation close to climatology, showing a tendency of decreasing towards the
LOMR, whereas the 2009 again shows a cold situation in AB, increasing towards the LOMR.
35