# A 70-year perspective on water-mass transformation in the Greenland Sea: From thermobaric to thermal convection

Anna-Marie Strehl, Kjetil Våge, Lars H Smedsrud, Thibaut Barreyre

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Anna-Marie Strehl, Kjetil Våge, Lars H Smedsrud, Thibaut Barreyre. A 70-year perspective on water-mass transformation in the Greenland Sea: From thermobaric to thermal convection. Progress in Oceanography, 2024, 227, pp.103304. 10.1016/j-pocean.2024.103304. hal-05402758

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# A 70-year perspective on water-mass transformation in the Greenland Sea: From thermobaric to thermal convection

Anna-Marie Strehl<sup>a,∗</sup>, Kjetil Våge<sup>a</sup>, Lars H. Smedsrud<sup>a,b</sup>, Thibaut Barreyre<sup>c,d</sup>

## ARTICLE INFO

**Keywords:**

Greenland Sea
Water-mass transformation

Deep convection

Stratification

a-ocean

Thermobaric effect

Atmosphere-ice-ocean interaction

Ice-edge retreat

Odden

## ABSTRACT

The hydrography of the central Greenland Sea was reconstructed from observations including bottle measurements, Conductivity/ Temperature/ Depth (CTD) measurements, and Argo floats for the period 1950 to 2020. Greenland Sea Deep Water was renewed during bottom-reaching convection prior to the mid-1980s, facilitated by the thermobaric effect. During a period of shallow convection in the late 1980s and early 1990s, a stratification maximum formed and isolated the deep from the intermediate Greenland Sea. As a consequence, convection was limited to depths shallower than 2000 m during the past decades and a new class of intermediate water formed instead of deep water. The initial cause for the formation of the stratification maximum was a near-surface freshwater anomaly. A subsequent, rapid temperature and salinity increase in the upper 2000 m resulted in an overall density reduction of the intermediate water which strengthened the stratification maximum. Along with the transition from formation of deep water to formation of intermediate water, the Greenland Sea became temperature-stratified at intermediate depths. This regime-shift in stratification can be traced to increased temperature and salinity in the inflowing Atlantic-origin Water. Below the stratification maximum, the Greenland Sea Deep Water became warmer and more saline, predominantly caused by lateral mixing with deep water masses from adjacent basins. The hydrographic changes in the Greenland Sea were investigated in the context of a reduction of the sea-ice extent and associated changes in winter heat loss. While interannual variability of convection depth may depend on atmospheric forcing, we found that the decadal variability of water-mass transformation in the Greenland Sea was largely determined by the hydrographic structure of the water column.

## 1. Introduction

The Nordic Seas are important for the supply of dense water to the lower limb of the Atlantic Meridional Overturning Circulation (AMOC) (Gebbie and Huybers, 2010; Chafik and Rossby, 2019). Warm and saline Atlantic Water flowing across the Greenland-Scotland Ridge into the Nordic Seas becomes colder and denser as it encircles the Nordic Seas and Arctic Ocean (Mauritzen, 1996) and is gradually transformed into Atlantic-origin Water. In the Iceland and Greenland Seas (Fig. 1), water-mass transformation leads to the formation of Arctic-origin Water, which is generally colder and denser than Atlantic-origin Water. Eventually, the transformed water masses return to the Greenland-Scotland Ridge, where they depart the Nordic Seas as over-flow plumes (Østerhus et al., 2019). Waters that are too deep to spill across the shallow ridge (with sill depths of 620 m in Denmark Strait and 850 m in Faroe Bank Channel, Blindheim, 1990, and Fig. 1) constitute

the deep waters of the Nordic Seas (Peterson and Rooth, 1976; Swift and Aagaard, 1981) and the Arctic Ocean (Langehaug and Falck, 2012). Beyond the contribution to the large-scale ocean circulation by the formation of dense water masses, convection also leads to efficient uptake of atmospheric gases such as carbon dioxide and oxygen (Davila et al., 2022) and facilitates vertical redistribution of nutrients within the ventilated part of the water column. Thus, deep and intermediate ocean convection is crucial for climate and ecosystems below the epipelagic zone.

The Greenland Sea and its importance to the overturning in the Nordic Seas have been investigated extensively since the early 20th century, when it was determined that the deep water masses of the Nordic Seas originate in the Greenland Sea (Helland-Hansen and Nansen,

<footnote><sup>a</sup> Geophysical Institute, University of Bergen and Bjerknes Centre for Climate Research, Bergen, Norway</footnote>

<footnote><sup>b</sup> University Centre in Svalbard, Longyearbyen, Svalbard and Jan Mayen</footnote>

<footnote><sup>c</sup> Geo-Ocean, CNRS, Univ Brest, Ifremer, UMR6538, F-29280 Plouzane, France</footnote>

<footnote><sup>d</sup> Department of Earth Science, Centre for Deep Sea Research, University of Bergen, Bergen, Norway</footnote>

<footnote>* Corresponding author at: Geophysical Institute, University of Bergen and Bjerknes Centre for Climate Research, Bergen, Norway. E-mail address: anna-marie.strehl@uib.no (A.-M. Strehl).</footnote>
1909). Numerous hydrographic and geochemical profiles obtained be-
tween then and the mid-1980s have documented bottom-reaching con-
vection (e.g., Meincke et al., 1992; Bönisch et al., 1997), but not
since. This abrupt and permanent transition from deep to intermediate
convection over only a few years provides a rare opportunity to achieve
a deeper understanding of the conditions leading to deep convection.
By combining all available data from the Greenland Sea since 1950
into a single, comprehensive data set, we aim to provide a consistent
history of the development of the hydrography and stratification in the
Greenland Sea over the past 70 years. Our focus is the transition from
deep- to intermediate-water production. We find that a stratification
maximum, which has separated the deep from the intermediate Green-
land Sea since the 1990s (Karstensen, 2005; Brakstad et al., 2019), has
been of major importance to this transition and examine the causes and
consequences of its formation. Furthermore, we explore the role of the
thermobaric effect for bottom-reaching convection. Sea-ice loss (e.g.,
Selyuzhenok et al., 2020) and ensuing changes in atmospheric forc-
ing (Moore et al., 2015) are considered in the second part of our
analysis. Finally, we discuss the role of water-mass transformation in
the Greenland Sea for the deep circulation of the Arctic Mediterranean
and the composition of the overflow water. To do justice to the com-
plexity of the topic and the extensive research that has been conducted
in the Greenland Sea, we start by reviewing the current understanding
and compare our results to previous findings in the discussion.

1.1. Water masses and circulation in the Nordic Seas

The Nordic Seas can be subdivided into three domains (Carmack and Aagaard, 1973; Swift and Aagaard, 1981; Swift, 1986). The Atlantic domain in the eastern part of the Nordic Seas is characterized by warm and saline Atlantic Water. Cooling of the Atlantic Water by heat loss to the atmosphere leads to transformation into Atlantic-origin Water. Cold and fresh surface water, Polar Water (PW), flows from the Arctic Ocean through Fram Strait into the Polar domain in the western part of the Nordic Seas (Foldvik et al., 1988). Due to the low surface salinity, which causes substantial near-surface stratification, water-mass transformation is limited here. The Arctic domain is situated between the Polar and Atlantic domains and is characterized by high near-surface density, low stratification, and large seasonal variability of surface temperature and salinity. In this domain, Arctic-origin Water is formed by water-mass transformation due to atmospheric cooling, mixing with subsurface Atlantic-origin Water and to a lesser extent entrainment of PW (Swift and Aagaard, 1981; Rudels et al., 2002; Svingen et al., 2023). The Arctic Front, which separates the Arctic domain from the Atlantic domain, follows the Knipovich, Mohn and Jan Mayen Ridges and is relatively stable (Raj et al., 2019), while the Polar Front, marking the boundary between the Polar and the Arctic domains, has a seasonal migration (Våge et al., 2018; Spall et al., 2021; Liu et al., 2022). During summer, the Polar domain extends farther east and a fresh surface layer covers much of the Greenland and Iceland Seas (Spall et al., 2021). In winter, onshore Ekman transport induced by strong northerly winds shifts the Polar Front toward Greenland, resulting in a higher salinity during winter which preconditions the central Greenland and Iceland Seas for the formation of Arctic-origin Water by open-ocean convection (Våge et al., 2018; Spall et al., 2021).

The Greenland Sea is confined by the Mohn and Knipovich Ridges to the east, the Greenland slope to the west, and the West Jan Mayen Ridge to the south (Fig. 1). On the eastern side of the Mohn and Knipovich Ridges, the western branch of the Norwegian Atlantic Current (the Norwegian Atlantic Frontal Current) transports Atlantic Water and Norwegian Sea Deep Water (NSDW) northward, while the East Greenland Current (EGC) carries PW, Atlantic-origin Water, and Arctic Ocean Deep Water (AODW) southward along the Greenland Slope (Fig. 1). Instabilities within these currents lead to the formation of eddies that exchange water masses between the central Greenland Sea and the boundary currents (Jeansson et al., 2008; Ypma et al., 2020; Spall

<figure><img src="image_2.png" /><figcaption>Fig. 1. Map of the Nordic Seas. The background color indicates bathymetry. The orange line highlights the 3000 m isobath defining the central Greenland Sea. The orange dot marks the center of the Greenland Sea Gyre (Moore et al., 2015; Brakstad et al., 2019). The currents surrounding the Greenland Sea are the Norwegian Atlantic Current (NAC), the East Greenland Current (EGC), and the Jan Mayen Current (JMC). Water masses transported by these currents are color-coded according to the legend in the top left corner of the figure. Exchange of deep water between the Greenland and Norwegian Seas through the Jan Mayen Channel is indicated by the yellow double-headed arrow. The green M marks the location of Ocean Weather Station M.</figcaption></figure>

et al., 2021), which brings Atlantic-origin Water from both the EGC and NwAFC into the central Greenland Sea. The Jan Mayen Current branches off the EGC north of the West Jan Mayen Ridge and transports PW and Atlantic-origin Water into the central Greenland Sea (Bourke et al., 1992; Pellichero et al., 2023). Another important pathway for the exchange of water masses is the 2200 m deep Jan Mayen Channel, which provides a deep connection between the Greenland Sea and the Norwegian Sea (Østerhus and Gammelsrød, 1999) (yellow arrow in Fig. 1).

A combination of several factors makes the Greenland Sea an ideal location for deep-water production. A cyclonic gyre leads to an upward doming of isopycnals in the center of the Greenland Sea (Clarke et al., 1986; Marshall and Schott, 1999). The magnitude of the doming varies, but lifting of up to 450 m has been documented (Marshall and Schott, 1999; Somavilla, 2019). This reduces stratification at intermediate depths and preconditions the central Greenland Sea for deep convection (Quadfasel and Meincke, 1987; Bogorodsky et al., 1993; Somavilla, 2019). The gyre circulation limits exchange with surrounding water masses and increases residence time in the central Greenland Sea (Straneo et al., 2003; Messias et al., 2008). This exposes the water masses inside the gyre to cooling over several winters and thereby contributes to the preconditioning for deep convection in subsequent winters (Helland-Hansen and Nansen, 1909; Aagaard et al., 1985; Clarke et al., 1986; Marshall and Schott, 1999). The Greenland Sea is regularly exposed to extreme cooling events such as cold-air outbreaks and polar lows (Papritz and Spengler, 2017; Terpstra et al., 2021). Cold-air outbreaks are responsible for 60–80% of winter heat loss in the Nordic Seas (Papritz and Spengler, 2017). The strong buoyancy loss during winter that is characteristic for the Greenland Sea (Marshall and Schott, 1999) further destabilizes the water column (Svingen et al., 2023; Renfrew et al., 2023). Surface cooling and freshwater removal due to sea-ice formation may lead to open-ocean convection if the water column becomes statically unstable. Due to the higher compressibility of cold than warm water, known as the thermobaric effect, very cold mixed layers can develop an internal
instability without further modification of the water's properties that
can lead to complete overturning of the water column (Akitomo, 1999).
This so-called thermobaric convection has been demonstrated in the
deep parts of the Weddel Sea (Akitomo, 1999), where the conditions
are similar to the central Greenland Sea.

## 1.2. Changes in the Greenland Sea water column

The coldest and densest waters in the Nordic Seas were formed
in the Greenland Sea (e.g., Helland-Hansen and Nansen, 1909; Pe-
terson and Rooth, 1976; Clarke et al., 1986; Dickson et al., 1996;
Jeansson et al., 2017; Brakstad et al., 2023). Bottom-reaching con-
vection regularly led to the renewal of Greenland Sea Deep Water
(GSDW) (Alekseev et al., 1994; Bönisch et al., 1997; Meincke et al.,
1997; Karstensen, 2005). As the deep water masses in the entire Nordic
Seas were nearly homogeneous (Swift, 1986; Blindheim, 1990), early
oceanographers concluded that the deep water spread from its origin
in the Greenland Sea into the rest of the Nordic Seas and Arctic
Ocean (Helland-Hansen and Nansen, 1909; Mosby, 1962). Although
deep convection occurred regularly before 1990 (Schlosser et al., 1991;
Bönisch and Schlosser, 1995; Bönisch et al., 1997), there were periods
of enhanced and reduced activity (Alekseev et al., 2001). Higher con-
vective activity in the late 1960s and early 1970s was associated with
low bottom-water temperatures (Meincke et al., 1992) and high oxygen
concentrations (Carmack and Aagaard, 1973; Clarke et al., 1990). The
salinity at deep and intermediate depths was relatively high, before a
sudden freshening occurred around 1980 (Clarke et al., 1986; Alekseev
et al., 2001).

During the 1980s, a remarkable transition took place: deep convection ceased and instead intermediate waters started to form (Karstensen, 2005). Since the mid-1990s, a stratification maximum has defined the base of the intermediate layer and prevented deeper convection (Brakstad et al., 2019). The deep-water properties in the central Greenland Sea were no longer determined by upper-ocean properties, but by mixing with other deep water masses such as AODW (Meincke et al., 1997; Somavilla et al., 2013) and NSDW (Østerhus and Gammelsrød, 1999). Since the cessation of deep convection, the GSDW has warmed (Budéus et al., 1998; Østerhus and Gammelsrød, 1999; Alekseev et al., 2001; Somavilla et al., 2013) and the salinity has increased (Bönisch et al., 1997; Budéus et al., 1998). Karstensen (2005) projected from the warming rate observed during the 1990s, that by approximately 2025 the deep Greenland Sea should match the properties of AODW unless deep convection occurs again. At intermediate depths, a temperature increase since the 1990s has been documented (Brakstad et al., 2019), while the salinity and oxygen concentration started to increase only after 2000 (Lauvset et al., 2018; Brakstad et al., 2019). The cause of the salinity increase in the Greenland Sea was traced to the Atlantic Water inflow into the Nordic Seas with a time-lag of 3 years (Lauvset et al., 2018). These property changes led to an overall density reduction of the intermediate water in the Greenland Sea and culminated in the formation of a new class of Greenland Sea Artic Intermediate Water (GSAIW) (Brakstad et al., 2019). During convection, excess heat and salt from the warmer and more saline Atlantic-origin Water that enters the Greenland Sea around 300 m depth (Svingen et al., 2023) are mixed into the GSAIW, such that GSAIW properties vary in accordance with the Atlantic-origin Water (Ronski and Budéus, 2005; Bashmachnikov et al., 2021).

The density reduction caused a redistribution of the deep and
intermediate waters originating in the Greenland Sea within the Nordic
Seas: waters formed within the Greenland Sea were previously too
dense to cross the Greenland-Scotland Ridge, but are now an important
component of the overflows (Huang et al., 2020; Brakstad et al., 2023).
The absence of deep-water renewal was associated with a reversal of
the flow in the Jan Mayen Channel from out of the Greenland Sea
to into the Greenland Sea in the 1990s, which has been found in
both observational (Østerhus and Gammelsrød, 1999) and numerical
studies (Whitehead, 2000; Kovalevsky and Bashmachnikov, 2020). The
supply of GSDW into the Eurasian Basin of the Arctic Ocean also ended
around this time (Langehaug and Falck, 2012).

## 1.3. Ice-edge retreat in the Greenland Sea

The remarkable transition from deep to intermediate convection
substantially altered the water column, but significant changes also
took place at the surface. During the 1960s and 1970s, sea ice regularly
extended across the entire Greenland Sea (Vinje, 1976). Characteristic
for the Greenland Sea during the 1980s and 1990s was the formation
of an ice tongue called Odden (Wadhams, 1981), which emerged above
the Jan Mayen Current and extended far into the central Greenland Sea
leaving only an open patch in the northwestern part of the Greenland
Sea called Nordbukta (Vinje, 1976; Shuchman et al., 1998; Wadhams,
1999; Comiso et al., 2001). Both local sea-ice formation and accumu-
lation of imported sea ice were found to be important contributors
to the appearance of the Odden (Comiso et al., 2001). During the
1990s, the ice edge retreated farther westward, leading to less frequent
appearances of Odden and some ice-free winters. Since 2000, the
Greenland Sea has largely been ice free (Kern et al., 2010; Selyuzhenok
et al., 2020).

The location of the ice edge plays an important role for deep convection. Two competing mechanisms take place in the marginal ice zone (Hibler and Bryan, 1987): on the one hand, low surface salinity, which is characteristic for this boundary region where sea-ice melt and growth take place, leads to a high surface stratification that inhibits convection. On the other hand, the marginal ice zone is where cold and dry polar air first encounters the relatively warm ocean. The large air-sea temperature differences and low humidity lead to enhanced sensible and latent heat fluxes from the ocean to the atmosphere (Marshall and Schott, 1999; Harden et al., 2015; Spensberger and Spengler, 2021). The formation of Odden/Nordbukta has been associated with enhanced convection in the Greenland Sea (Roach et al., 1993; Schott et al., 1993). Due to the ice-edge retreat of the recent decades, the Greenland Sea Gyre is increasingly far removed from the ice edge and thus from the location of the largest turbulent heat fluxes. The result is a reduction in wintertime heat loss from the gyre of about 20% from 1997 to 2014 (Moore et al., 2015).

## 2. Data and methods

### 2.1. Hydrographic data

#### Central Greenland Sea

We use a data set consisting of hydrographic data from various
sources covering the entire Nordic Seas from 1950 to 2020 (Brak-
stad, 2023b). It contains a combination of shipboard Conductivity/
Temperature/ Depth (CTD) profiles, eXpendable Conductivity/ Tem-
perature/ Depth (xCTD) profiles, bottle measurements, and data from
Argo floats. The older data (before 2000) are predominantly shipboard
measurements, while the last 20 years are dominated by profiles from
Argo floats (Brakstad et al., 2023). The data set has been augmented
with measurements from the *Climatological Atlas of the Nordic Seas and
Northern North Atlantic* (Korablev et al., 2014), various hydrographic
surveys (Holfort, 2001; Schauer, 2010; Logemann, 2018; Quadfasel,
2018), and data from a profiling mooring deployed in the central
Greenland Sea from summer 1999 to summer 2009 (Budéus et al.,
2020) that were not included in the original data set. A complete list of
all included data sets can be found in Table A1.1. Details of the filtering,
gridding, and quality control of the hydrographic data set are provided
in Appendix A2.

#### Adjacent basins

Measurements obtained at Ocean Weather Station M (Fig. 1) at
2000 m depth are representative of the hydrographic properties of
NSDW (Gammelsrød et al., 1992; Østerhus and Gammelsrød, 1999;
Dickson et al., 1996). As the publicly available data from Station M
have not been calibrated, we extracted measurements in the vicin-
ity of Station M between 1900 and 2100 m depth located between
65.5 and 67.0°N, and 1 and 3°W from the hydrographic data set published and quality controlled by Brakstad (2023b). Profiles representing the properties of AODW, which enters the Nordic Seas via Fram Strait, were extracted from the same hydrographic data set within a box in western Fram Strait, reaching from 78.8 to 78.9°N, and 3.8 to 0°W. Only measurements between 1900 and 2100 m depths were included in the analysis. CTD data obtained during annual monitoring surveys across Fram Strait were included in the Nordic Seas data set until 2014, data from the following years were added manually [data sets by Dodd et al. in the reference list].

## Uncertainties

The error estimates below are provided in original units. For our analysis, we apply the TEOS-10 standard (IOC et al., 2010). Absolute Salinity [g/kg] and Conservative Temperature [°C] are consistently used throughout the paper and are hereafter referred to as salinity and temperature.

The accuracy of the hydrographic measurements have improved substantially over time and are summarized in Table A.2. The first oceanographic pressure measurements were based on pressure-dependent expansion of crystals (Mohn, 1887). The technology has not changed fundamentally since then, but the accuracy improved significantly when digital measurements became possible. Modern quartz crystal resonators have an accuracy of ±0.015% of full scale range (Driemel et al., 2017), which corresponds to 0.6 dbar for sensors with a pressure rating of 4000 dbar. This is about one order of magnitude improvement over the past 150 years (Mohn, 1887).

Before the 1970s, salinity was measured by titration of water samples. This method has an accuracy of ±0.01 ‰, but constant offsets of up to ±0.03 ‰ may occur (Helland-Hansen and Nansen, 1909). The Practical Salinity Unit (psu) was introduced when titration was replaced by conductivity measurements and meant to reflect ‰ (Dauphinee, 1980). Since the late 1970s, salinity has routinely been estimated from the conductivity of sea water with an accuracy of ±0.002 psu (Dickson et al., 1996). Modern CTDs achieve an accuracy of ±0.001 psu (Ullman and Hebert, 2014) to ±0.003 psu (Krahmann et al., 2009). For xCTDs, the uncertainties were an order of magnitude larger in the mid-1990s (±0.05 psu, Mizuno and Watanabe, 1998), but improved to approximately CTD quality over the recent decades (Renfrew et al., 2023). In the typical salinity ranges of the Nordic Seas (27-36 g/kg), the relationship between salinity as a unitless quantity (psu) and salinity in g/kg is approximately linear and can be estimated with $S_{g/kg} = 1.0048 \cdot S_{psu}$.
The difference in accuracy is on the order of $10^{-6}$.

In the early 20th century, temperature measurements were obtained with an accuracy ranging from ±0.01 °C to ±0.05 °C, depending on the thermometer used (Helland-Hansen and Nansen, 1909). While reversing thermometers were used until the early 1980s, the accuracy had improved to about ±0.01 °C to ±0.02 °C by the 1950s (Aagaard, 1968), where it remained until the 1990s (Dickson et al., 1996). For current CTD temperature measurements, typical accuracies are ±0.001 to 0.002 °C (Krahmann et al., 2009; Ullman and Hebert, 2014), but higher accuracies can be obtained with careful calibration (e.g., Uchida et al., 2007). The accuracy of temperature measurements made by xCTD ranges from 0.05 °C in the mid-1990s (Mizuno and Watanabe, 1998) to 0.001 °C in the late 2010s (Renfrew et al., 2023).

As Argo floats are not collected at the end of their batteries' lifetimes, sensor drift needs to be corrected using statistical methods and climatological data instead of comparison with externally calibrated measurements (Wong et al., 2020). Argo data corrected in this way after completion of the float's mission are referred to as delayed-mode data and are the only Argo data source used for this study. The accuracies of the Argo measurements are estimated to ±0.002 °C for temperature and ±0.01 g/kg for salinity (Wong et al., 2020).

## Distribution of the observations in space and time

While most hydrographic surveys in the Greenland Sea took place in summer, Argo data are collected autonomously throughout the year.

<figure><img src="image_3.png" /><figcaption>**Fig. 2.** Seasonal data distribution. For each of the three periods indicated in the legend, the total number of profiles from each month is shown by the colored bars. The median number of profiles per month for each period is indicated by the colored lines.</figcaption></figure>

As the Argo program has only been operative since the early 2000s, there are more summer than winter profiles from the early period. Our hydrographic data set therefore has a summer bias (Fig. 2), which is detrimental for the investigation of processes happening on seasonal timescales, such as convection. We have accounted for this in the analysis by applying different temporal resolutions during the gridding of the data set (for more details see Appendix A2) and by introducing the σ-π-distance as a measure for vertical homogeneity (Section 2.5), which facilitates reconstruction of winter convection depth from summer data (Section 3.1).

The increase in number of profiles in the third period (Fig. 2) is due in large part to the inception of the Argo program. However, Argo floats only sample the upper 2000 m depth. Below this depth, only sparse observations exist after 2010 (illustrated in the bottom row of Fig. 4). Below 2000 m, particularly after the cessation of bottom-reaching convection, variability is low and has a multi-year time scale. Changes occurring over pentads or decades will therefore be adequately represented by the nearly annually available data. This is reflected in our gridding procedure (Appendix A2).

The number of profiles per time interval has increased substantially, though not continuously, over the 70-year period. For instance, with the advent of the Argo program in 2001, the sampling rate more than doubled compared to the previous decades. To ensure a consistent coverage, we identified three periods which are characterized by a relatively uniform data density: 1950–1972, 1973–2000, and 2001–2019 (more details are provided in Appendix A2 and Fig. A2.1).

Before 1973, the measurements were sparse in time and space and the vertical resolution was low (the average vertical resolution was 120 m). For the period 1973–2001, when CTD profiles started to complement bottle measurements, the mean vertical resolution increased substantially to 3 m, and further to 1.6 m after 2001 (Fig. A2.1). Until the advent of the Argo program in 2001, the profiles were mainly distributed along repeat hydrographic sections. The first period, 1950–1972, is dominated by data along a section at 74.5°N across the central Greenland Sea (Fig. 3). Despite an accumulation of data along sections in the first two periods, profiles from everywhere within the contour are available in all three periods, such that the entire central Greenland Sea is represented throughout the 70-year period.
<figure><img src="image_4.png" /><figcaption>Fig. 3. Locations of profiles in the central Greenland Sea, color-coded according to when they were obtained. For each period the total number of profiles is given in the lower left corner. The thin lines are the 1000 m (light gray), 2000 m (blue), and 3000 m (black) isobaths. The thick black line outlines the central Greenland Sea defined by the 3000 m isobath following Latarius and Quadfasel (2016). The thick gray line outlines the Greenland Sea Gyre defined by a closed contour of dynamic topography (Moore et al., 2015).</figcaption></figure>

## 2.2. Heat-flux and sea-ice data

We used the ERA5 reanalysis (Hersbach et al., 2020) to investigate ocean-atmosphere heat exchange and sea-ice concentration. The reanalysis is produced by a four-dimensional data assimilation into the ECMWF Integrated Forecasting System (IFS). The currently available product was calculated with time steps of about 12 min on a grid with a spatial resolution of 31 km (Hersbach et al., 2020). For latent and sensible heat fluxes and sea-ice concentration, hourly data are provided from January 1950 onward. For our analysis we averaged 24 hourly snapshots to obtain daily mean values.

Sea-ice concentration in ERA5 is based on available sea-ice products incorporating both satellite and in-situ data (Titchner and Rayner, 2014; Donlon et al., 2012; Good et al., 2020; Walsh et al., 2017). In the IFS, sea-ice melt and formation are predicted. For the reanalysis product, sea-ice observations are assimilated into the IFS and impact the interaction between atmosphere and ocean. It is thus advantageous to use both heat-flux and sea-ice fields from ERA5 to investigate atmosphere-ice-ocean interactions.

While ERA5 performs well over the open ocean, the sea-ice distribution in the marginal ice zone is overly smooth, resulting in an inaccurate representation of the heat fluxes in this region (Renfrew et al., 2021). We therefore consider heat fluxes only from the ice-free portion of the Greenland Sea. As the presence of sea ice limits interaction between the ocean and the atmosphere, the contribution from the marginal ice zone is reduced compared to the open ocean. Hence, the interannual variability is well represented by excluding the marginal ice zone from the calculations of both extreme and total heat loss.

## 2.3. Geothermal measurements

To assess the impact of geothermal heating on the development of the GSDW, we extracted measurements of sea-floor heat flow from the Global Heat Flow Database (GHFDB, Fuchs et al., 2021). As an assessment of uncertainty, the values obtained for the central Greenland Sea were compared with global averages for regions with a similar crustal age. The crustal age stems from the global data set of present-day oceanic crustal age (Müller et al., 2019; Seton et al., 2020).

## 2.4. Distinction between Greenland Sea intermediate and deep water masses

Following Rudels et al. (2002), the isopycnal distinguishing intermediate and deep water masses is $\sigma_{0.5} = 30.444 \frac{kg}{m^3}$, using the EOS-80 standard (IOC et al., 2010). Density calculated using the new TEOS-10 standard is slightly higher and the differences are non-linear (Fig. A3.1). At $\sigma_{0.5, EOS-80} = 30.444 \frac{kg}{m^3}$, the TEOS-10 standard potential density is on average $0.004 \frac{kg}{m^3}$ higher. We therefore use an adjusted value of $\sigma_{0.5, TEOS-10} = 30.448 \frac{kg}{m^3}$ to distinguish intermediate and deep water masses.

## 2.5. $\sigma-\pi$ distance

Potential spicity, $\pi$, is a variable defined to be orthogonal to density, $\sigma$ (Huang et al., 2018). T.S properties of water masses can be transformed into $\sigma-\pi$ properties without loss of information. In $\sigma-\pi$ space, meaningful distances between water masses can be calculated due to the orthogonality of the coordinates. The $\sigma-\pi$ distance $D$ between two water masses with properties ($\sigma_1, \pi_1$) and ($\sigma_2, \pi_2$) can be defined in a Euclidean way (Huang et al., 2018):

$$D = \sqrt{(\sigma_1 - \sigma_2)^2 + (\pi_1 - \pi_2)^2}$$

We interpret the $\sigma-\pi$ distance as a measure of similarity between water masses. The smaller the distance $D$, the more alike are the water masses.

## 2.6. Determination of mixed-layer depth

Late-winter mixed layers were determined by a method that involves visual inspection of each profile. We define late winter as February to April, but due to the scarcity of winter measurements before 1989 we also considered profiles from May, when remnants of convection are often evident. The mixed-layer depth was first suggested by two automated routines. One of the routines estimates the mixed-layer depth by applying a density-difference criterion relative to the surface (Nilsen and Falck, 2006): the shallowest depth where the density increased by more than $\Delta\rho$ that corresponds to a temperature reduction by $0.8^\circ C$ is associated with the mixed-layer depth. To better match the conditions in the central Greenland Sea, where low surface temperatures are prevalent throughout the year, we reduced the threshold value to $0.2^\circ C$. The other routine identifies the base of the mixed layer as the shallowest extreme in the curvature of the temperature profile (Lorbacher et al., 2006). For each late-winter profile, the automatically detected mixed-layer depth was visually inspected and corrected (Brakstad et al., 2019), if necessary, by the following method: if the variability within the mixed layer exceeded two standard deviations, the mixed-layer depth was adjusted manually until the variability of the mixed layer was within two standard deviations (as suggested by Pickart et al., 2002). Mixed-layer depths between 1990 and 2016 were determined by Brakstad et al. (2019) following the same procedure.

While mixed layers are usually determined by the deepest level where temperature, salinity, and density are constant from the surface, we additionally considered mixed layers that were isolated from the surface as evidence of recent convection. After active convection ends, the surface restratifies rapidly and thereby shields the mixed layer from further interaction with the atmosphere (Jones and Marshall, 1997). As such, the isolated mixed layer remains detectable until it is eroded by lateral mixing. As the Greenland Sea was characterized by a weak background stratification during most of the 70-year period, only nearly homogeneous profiles can be interpreted as remnants of convection. To ensure a high standard for the correct identification of such nearly homogeneous sub-surface mixed layers, the stratification of each late-winter profile was visually compared to a February 1971 profile that shows evidence of bottom-reaching convection below a
shallow surface mixed-layer of less than 200 m depth (Malmberg, 1983, the reference profile is shown in Fig. A3.1). For comparison, the profiles in our data set showing evidence of recent bottom-reaching convection are displayed in Fig. A4.2.

## 2.7. Determination of thermobaric depth

The central Greenland Sea is sufficiently cold and deep that the thermobaric effect must be considered (e.g., Garwood et al., 1994; Akitomo, 2011). It stems from the non-linearity of the temperature-dependent thermal expansion coefficient: cold water is more compressible than warm water. As a consequence, cold water increases more rapidly in density with depth than warmer water. This effect is very weak and only important below the so-called thermobaric depth. Beneath this depth, stratification caused by temperature depends on both the temperature gradient and the compressibility of sea water. To determine the thermobaric depth, the thermal expansion coefficient $\alpha$ is assumed to have a constant contribution $\alpha_0$ and to increase linearly with depth $z$ where $\alpha_1$ is a linear coefficient (Akitomo, 1999).

$$ \alpha = \alpha_0 + \alpha_1 \cdot z $$

The depth where the linear term first dominates the constant term is defined as the thermobaric depth. This leads to the following expression for the thermobaric depth $H_a$:

$$ H_a \geq \frac{\alpha_0}{\alpha_1} $$

The thermobaric depth may indicate whether individual episodes of convection were enhanced by the thermobaric effect.

## 2.8. Stratification

Stratification is expressed using the buoyancy frequency $N^2$. As the thermobaric effect affects stratification, we accounted for the depth-dependence of the density $\rho$ when computing the buoyancy frequency. Density as well as density change were referenced to the pressure at depth $z$ instead of surface pressure. This results in the following formulation for the buoyancy frequency $N^2$, where $g$ is the gravitational acceleration:

$$ N^2(z) = g \frac{1}{\rho_z} \frac{\partial \rho}{\partial z} \Big|_z $$

## 2.9. Classification of stratification into $\alpha$- and $\beta$-oceans

Stratification $N^2$ can be separated into contributions from temperature ($N_T^2 = \alpha \frac{dT}{dz}$) and salinity ($N_S^2 = \beta \frac{dS}{dz}$) (Carmack, 2007). $\alpha$ and $\beta$ are the thermal expansion and haline contraction coefficients, respectively. Both contributions can be positive (stable) or negative (unstable). A temperature-sustained stratification ($N_T^2 > 0, N_S^2 < 0$) is referred to as an $\alpha$-ocean, salinity-sustained stratification is called a $\beta$-ocean (Carmack, 2007). This distinction is not defined when both contributions have the same sign, which has later been termed transition-ocean (e.g., Stewart and Haine, 2016). To classify the stratification unambiguously into $\alpha$, $\beta$, or transition-ocean, we follow the definition of Stewart and Haine (2016), which is based on the spice frequency $K^2$. Spice frequency may be understood as the counterpart of buoyancy frequency, indicating the strength of the along-isopycnal gradient instead of the diapycnal gradient (stratification). Generally, positive $K^2$ is associated with $\alpha$-oceans, while negative $K^2$ characterizes $\beta$-oceans. If, however, the variability of $K^2$ is greater than its absolute value, the tendency toward $\alpha$-$\beta$-stratification is insignificant, and the water column is considered a transition-ocean. The mathematical formulation follows Stewart and Haine (2016):

$$ \alpha\text{-ocean}: \quad \overline{K^2} - \sigma_{K^2} > 0 $$

$$ \beta\text{-ocean}: \quad \overline{K^2} + \sigma_{K^2} < 0 $$

$$ \text{transition-ocean}: \quad |\overline{K^2}| - \sigma_{K^2} \leq 0 $$

$$ \text{with} \quad K^2 = g \cdot \left( \alpha \frac{\partial T}{\partial z} + \beta \frac{\partial S}{\partial z} \right) $$

$$ \text{and} \quad \sigma_{K^2} = \sqrt{(K^2)^2 - \overline{K^2} \overline{K^2}} $$

## 2.10. Pattern classification

Although Odden was a seemingly unambiguous sea-ice feature in the central Greenland Sea, the high temporal and spatial variability in sea-ice concentration leads to a large variety of possible realizations recognizable as Odden. To find patterns in the spatial variability of sea-ice concentration, we used a technique known as a Kohonen map, also called a self-organizing map (Kohonen, 1982). The pattern-recognition algorithm iteratively sorts vectors into a pre-determined number of patterns. Each vector represents a specific state of the system, in our case the spatial distribution of sea-ice concentration at each time step. Initially, the vectors are randomly attributed to a pattern. Then, the distance of each vector to each pattern is calculated as the average Euclidean distance to the vectors contained in each pattern. The vectors are then associated with their closest pattern, respectively. In this way, the vectors are sorted into the patterns by minimizing the least-square variability within each pattern. The optimization takes place over a given number of steps, so-called epochs. In our case, convergence was achieved after 80 to 100 epochs. To ensure full convergence, we used 150 epochs.

# 3. Long-term hydrographic development

The mid-1980s was a turning point for the Greenland Sea hydrography. The cessation of bottom-reaching convection set the developments of the deep and intermediate water masses on different trajectories.

## 3.1. 1950–1985: Formation of Greenland Sea Deep Water

Between 1950 and 1985, the central Greenland Sea below 500 m was remarkably vertically homogeneous: changes in temperature and salinity occurred for the most part simultaneously over the entire water column (Fig. 4). Two phases stand out of which the first was characterized by particularly low temperature lasting from 1965 to 1985 (Fig. 4a), and the second by low salinity between 1980 and 1990 (Fig. 4b). Over the entire 35-year period, isopycnal depths fluctuated by up to 1000 m, but there were no consistent trends (Fig. 4c).

To estimate the degree of vertical mixing within the water column, we computed the $\sigma$-$\pi$ distance (Section 2.5) between the hydrographic properties near 750 m and the rest of the water column (Fig. 4d). In the absence of convection, the water column below the mixed layer restratifies by mixing with water masses originating outside the central Greenland Sea and the $\sigma$-$\pi$ distance increases. A vertically homogeneous water mass with uniform properties and low $\sigma$-$\pi$ distance indicates recent convection. The depth of 750 m was chosen because it is well below the strongly stratified layer that forms during summer (dark green layer in the upper ocean in Fig. 4e), and represents the water mass that formed or was present in the previous winter. Due to low temporal resolution and lack of wintertime measurements in the early period (Fig. 2), the high $\sigma$-$\pi$ distances in the upper water column have a summer bias. There, wintertime convection can only be discerned in the high-resolution period after 2001 (Fig. 4d). Below the seasonally-varying upper layer, changes occur more slowly and are better suited for application of the $\sigma$-$\pi$ distance metric to assess mixed-layer depth. Coinciding with the low-temperature phase between 1965 and 1985, the hydrography of the Greenland Sea was at times nearly uniform from 750 m depth to the bottom (Fig. 4d). This was likely caused by bottom-reaching convection that homogenized the entire water column. Between periods of convection to the bottom in the 1970s, higher values of $\sigma$-$\pi$ distance appeared around 2500 m
<figure><img src="image_5.png" /><figcaption>**Fig. 4.** Development of (a) temperature, (b) salinity, (c) density, (d) σ-π distance to the mean properties across the 700 - 800 m depth range, (e) buoyancy frequency, and (f) classification of stratification into α- and β-ocean over the 70-year period. Note the non-linear color bars. The black dots in (d) are mixed-layer depths estimated from each winter profile. The black dots and line in (f) are the thermobaric depth $H_w$ calculated for each profile and the linear fit for the period 2001–2020, respectively. Density contours in (a)–(e) are referenced to 3000 m except the thicker line, which is referenced to 500 m and marks the isopycnal distinguishing intermediate and deep water masses. Each black line at the top indicates the presence of a profile, red and white colors indicate summer (May-Oct) and winter (Nov-Apr) months, respectively. At the bottom, profiles extending deeper than 2000 m are marked.</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 4. (continued).</figcaption></figure>
(Fig. 4d). Local maxima in $\sigma-\pi$ distance indicate that the hydrographic properties of the water column were changed at that depth with limited modification of the water mass above or below. As the water column remained statically stable, the intrusion must have occurred isopically. The $\sigma-\pi$ distance maxima localized at around 2500 m depth in the late 1970s and 1980s thus suggest that lateral inflow was enhanced at this depth. Considering the evolution in $\sigma-\pi$ distance, it appears that the most vigorous convection occurred between 1965 and 1970, followed by a period of shallower convection until the late 1970s, before bottom-reaching convection again nearly homogenized the entire water column. After the cessation of bottom-reaching convection around 1985, the hydrographic properties of the intermediate and deep water masses in the Greenland Sea have become increasingly distinct (Fig. 4d).

Consistent with bottom-reaching convection regularly occurring between 1950 and 1985, the water column below 750 m was very weakly stratified (Fig. 4e). The near-surface seasonal cycle of weak stratification during winter and strong stratification in summer can be discerned once the temporal resolution increases toward the end of the period (Fig. 4e). Before 2001, the seemingly constant high surface stratification only represents summertime conditions. Haline stratification dominated the fresh upper layer (Fig. 4b) due to a pronounced vertical salinity gradient, consistent with classification as $\beta$-ocean (Section 2.9, Fig. 4f). At greater depth, the vertical temperature and salinity gradients contributed equally to the weak stratification, such that this part of the water column can be classified as transition-ocean (Fig. 4f). Due to the cold temperatures, the thermobaric depth (Section 2.7) was relatively shallow and stable at approximately 1500 m (Fig. 4f). The deepest part of the water column was temperature-stratified (Fig. 4f), which can primarily be ascribed to the thermobaric effect.

## 3.2. 1985–2020: Formation of Greenland Sea Arctic Intermediate Water

Between 1985 and 2020, the density of the entire water column decreased (Fig. 4c). While the salinity increased (Fig. 4b), the magnitude of the increase was not sufficient to overcome a general warming trend (Fig. 4a). The density decrease led to a descent of the 30.448 $\frac{kg}{m^3}$ isopycnal, which distinguishes deep from intermediate water masses (Section 2.4), implying that the volume of GSDW has been substantially reduced. In the 1990s, a stratification maximum emerged around 1000 m, which marked the top of the GSDW layer (Fig. 4e). The emergence of the stratification maximum coincided with the cessation of bottom-reaching convection. Above the stratification maximum, a new class of less dense intermediate water, GSAIW, started forming (Brakstad et al., 2019).

While the deep central Greenland Sea below 750 m was for the most part well mixed prior to 1985, the stratification maximum has since kept the intermediate and the deep water masses apart. The $\sigma-\pi$ distance between water properties at 750 m and the abyss gradually increased over the last 40 years of the record (Fig. 4d), implying that GSAIW and GSDW have become increasingly distinct and that negligible bottom-reaching convection or mixing between the two water masses has taken place. During the 1990s, the stratification maximum descended to a depth of around 2000 m, where it has remained since the beginning of the 2000s and become increasingly prominent (Fig. 4e).

The temperature and salinity of the GSDW increased continuously over the last 40 years (Figs. 4a and b). By contrast, the GSAIW properties remained nearly constant during the 1990s, until its temperature and salinity also started increasing around 2000 (Figs. 4a and b).

Seasonal salinity maxima regularly occurred in the upper 200-500 m, but wintertime convection distributed the high-salinity water across the entire GSAIW layer and led to its overall salinity increase (Fig. 4b). At this depth range, Atlantic-origin Water spreads into the Greenland Sea (Svingen et al., 2023), and the temperature and salinity increase of the GSAIW can likely be traced to changes in the properties of the inflowing Atlantic-origin Water (e.g., Alekseev et al., 2001;

Lauvset et al., 2018) or increasing amounts of Atlantic-origin Water entering the Greenland Sea due to enhanced recirculation in Fram Strait (e.g., Muilwijk et al., 2018; Selyuzhenok et al., 2020; Bashmachnikov et al., 2021; Wefing et al., 2022). The impact of Atlantic-origin Water on GSAIW variability is further discussed in Section 3.4.

Below the stratification maximum, the salinity of the GSDW began to increase in 1990 at around 2500 m (Fig. 4b), which corresponds to the depth where enhanced lateral mixing occurred before 1985 (Fig. 4d). We may conclude that deep inflows from adjacent basins predominantly take place around 2500 m depth, which is consistent with other deep circulation schemes (Aagaard et al., 1985).

As the central Greenland Sea warmed, the thermobaric depth increased from a nearly constant depth near 1500 m to its current depth of approximately 2500 m (black dots and regression line in Fig. 4f). The descent began shortly after 2000, and since then the thermobaric depth has always been situated below the stratification maximum. The implications will be discussed in Section 3.3.

Below 2000 m, the stratification was temperature-dominated ($\alpha$-ocean) throughout the entire 70-year period. When GSDW was present between 500 and 2000 m, this depth-range was a transition-ocean. The stratification maximum that formed in the 1990s was salinity-stratified ($\beta$-ocean), indicating that its origin was a near-surface fresh anomaly. In the late 2000s, the stratification maximum became temperature-stratified. The GSAIW that formed above the stratification maximum after 2000 coincided with a transition toward an $\alpha$-ocean over the intermediate depth range. As a result of this development, the entire Greenland Sea has been temperature-stratified since the beginning of the 2010s.

## 3.3. Convection

We identified several very deep mixed layers revealing bottom-reaching convection between 1965 and 1985 (Fig. 5). Since 1985, mixed layers extending deeper than 1600 m have not been recorded. Due to the limited data coverage in the Greenland Sea, particularly during winter, the mixed-layer depth estimates are not sufficient evidence that convection ceased. Indirect methods drawing on both winter and summer data provide additional support for the assessment of mixed-layer depth. The $\sigma-\pi$ distance of the deep Greenland Sea to 750 m depth, based on both summer and winter profiles, is a measure of the water column's homogeneity. The good agreement between the mixed-layer depths and low $\sigma-\pi$ distances (Fig. 4d) indicates that the ventilation depth in the Greenland Sea has been accurately determined.

After the cessation of bottom-reaching convection in 1985, a period of shallow convection with mixed layers no deeper than 1000 m deep followed. In the following decade, along with the descent of the stratification maximum, the mixed-layer depths gradually increased. Since 2011, the mixed-layer depths have shoaled (Fig. 4d and 5) and the stratification maximum has intensified (Fig. 4e).

Based on the mixed-layer depth record (Fig. 5), we suggest that the long-term hydrographic development in the central Greenland Sea can be divided into three distinct phases. The first period, between 1965 and 1985, was characterized by bottom-reaching convection and a homogeneous water column. After the cessation of bottom-reaching convection in 1985, a period of shallow convection ensued and the stratification maximum emerged. In the final period, after 2000, convection to intermediate depth took place while the stratification maximum descended to its present depth near 2000 m and fortified. We emphasize that convection was not uniform during these periods, winters with convection shallower than 500 m occurred in all three phases (Fig. 5), but the maximum mixed-layer depth was distinct for each period.

Two classes of mixed layers are discernible: mixed layers shallower than 2000 m or deeper than 3000 m (Fig. 6). This distribution indicates that bottom-reaching convection ensues once a threshold depth has
<figure><img src="image_7.png" /><figcaption>Fig. 5. Development of mixed-layer depth over the 70-year period. The red dots show the maximum mixed-layer depth for each winter (January to May), the red errorbars correspond to half of the vertical resolution of the corresponding profile. The gray bars show the 80th percentile of all identified mixed-layer depths, shown by the black dots, each winter. The blue shading indicates typical convection depths for different periods within the 70-year record based on a qualitative assessment combining the observed mixed-layer depths and the $σ-π$ distance to 750 m.</figcaption></figure>

been reached. The very low temperatures ($< -1.2^\circ C$) of the mixed-layers extending to the bottom suggest that thermobaric convection has occurred. Sufficiently cold mixed layers exceeding the thermobaric depth will experience a faster density increase with further deepening than the ambient water. The higher density promotes further deepening of the mixed layer until a layer of equal or higher density is reached. In the weakly-stratified Greenland Sea, there are no such higher-density layers that could arrest the thermobaric sinking. Consequently, thermobaric convection would lead to bottom-reaching convection and a complete overturning of the water column. This hypothesis is corroborated by the fact that none of the mixed layers shallower than 2000 m reached their thermobaric depths (Fig. 6a). Due to the very weak stratification at depth (Fig. 4e), thermobaric convection may not have been a necessary condition for complete overturning, but the mechanism at least facilitates bottom-reaching convection once the thermobaric depth has been reached.

Coincident with the warming of the Greenland Sea, the thermobaric depth has increased (Fig. 4f). Since the stratification maximum formed, the thermobaric depth has consistently been located below this layer. To exceed the thermobaric depth and enter the thermobaric-convection regime, convection first has to overcome the stratification maximum. As such, the stratification maximum acts as a barrier for deep convection in general and for thermobaric convection in particular.

While the properties of bottom-reaching mixed layers are contained within a narrow range in both temperature and salinity, intermediate convection occurs within a much wider temperature range (Fig. 6). The salinity of mixed layers deeper than 500 m is at minimum 35.00 g/kg and mixed layers deeper than 1000 m have a salinity of at least 35.04 g/kg, suggesting that a salinity threshold needs to be overcome to achieve intermediate convection. This is consistent with previous findings that a minimum surface salinity of 34.88 g/kg (34.71 psu) in summer is necessary to achieve convection beyond 500 m (Dickson et al., 1988; Malmberg and Jónsson, 1997; Brakstad et al., 2019). While the salinity threshold for 1000 m-deep convection suffices also for bottom-reaching convection, the mixed-layer temperature decreases with increasingly deep convection. We conclude that salinity determines whether convection may occur in the central Greenland Sea, but that the final depth of convection depends on temperature. Convection is thus limited by

salinity but thermally driven. While this interpretation holds for the entire 70-year period (Fig. 6 contains mixed-layers from the entire 70-year period), we emphasize that it is particularly true for the most recent 20 years, when the stratification became temperature-dominated (Fig. 4f).

### 3.4. Development and modification of water masses in the Greenland sea

#### Development of mixed-layer properties

We assess the long-term hydrographic development of the densest waters in the Greenland Sea by calculating the mean hydrographic properties of the 20% deepest mixed layers from each winter. These values are then grouped into pentads to minimize interannual variability. Before 1985, longer periods were chosen such that the mean values represent approximately as many mixed layers as the five-year periods after 1985 (Fig. 7). The densest mixed layers, with properties similar to the GSDW, were formed before 1975 (Fig. 7). In the following decade, the mixed layers were generally colder, fresher, and less dense. This was partly due to the more frequent occurrence of winters with shallow convection only, but also the subset of mixed layers reaching the bottom were colder and fresher. This freshening of the Greenland Sea in the 1980s lagged by half a decade a fresh anomaly in the Faroe-Shetland Channel which propagated northward along the Norwegian slope during the late 1970s (Holliday et al., 2008). The origin of Greenland Sea freshening will be further elaborated on in the discussion.

During the period when only shallow ($< 1000 \ \text{m}$) convection took place (1985–1995) (Fig. 5), the salinity of the mixed layer was so low that even cooling to the freezing point would not have increased the density sufficiently to form GSDW (Fig. 7). In the late 1990s, the salinity started to increase. This coincided with increasing mixed-layer depths. By 2000, the mixed-layer salinity had reverted to values comparable to those when bottom-reaching convection took place, but a concurrent mixed-layer temperature increase of about $1/4^\circ C$ prevented formation of sufficiently dense water to renew the GSDW (Fig. 7). This marks the transition from the formation of GSDW to the formation of GSAIW in the central Greenland Sea (Brakstad et al., 2019). Over the following decades, the mixed-layer temperature and salinity continued to increase, while the density decreased. The mixed-layer salinity and temperature are currently at their highest within the 70-year record.
<figure><img src="image_8.png" /><figcaption>**Fig. 6.** Distribution of mixed-layer depths and corresponding hydrographic properties. The colors in (a) indicate the distance from the bottom of the mixed layer to the thermobaric depth. The thick black line marks the average thermobaric depth for the 70-year period and the thin black lines mark the standard deviation envelope. The gray area indicates the range of thermobaric depths. The statistical properties of the thermobaric depth are based on annual mean values, while the distance of the mixed-layer base to the thermobaric depth was calculated for each profile. The colors in (b) indicate the number of profiles contributing to each depth-salinity grid point.</figcaption></figure>

<figure><img src="image_9.png" /><figcaption>**Fig. 7.** Development of the hydrographic properties of the deepest 20% mixed layers over the 70-year period. The large circles represent the mean hydrographic properties of the mixed layers for the indicated periods. The size of each circle is scaled by the number of mixed layers contributing to the average, while the colors of the interior and perimeter indicate the average and maximum mixed-layer depths, respectively. The empty circles in the top left corner serve as legend and the numbers below indicate the number of mixed layers per data point. The light-blue errorbars mark the standard deviations of the hydrographic properties. The dark-blue dots reflect the evolution of the GSDW, which is shown in more detail in Fig. 10. The light-gray isopycnals are referenced to the surface, while the $σ_{0.5} = 30.448 \text{ kg/m}^3$ isopycnal separating intermediate and deep water masses is referenced to 500 m.</figcaption></figure>

## The impact of Atlantic-origin Water on mixed-layer salinity

During summer, when there is no active convection, the mixed layer restratifies. In the central Greenland Sea, restratification is most pronounced in the upper 50 m (Fig. 8a and b). While the surface layer freshens during summer, a sub-surface salinity maximum develops between 50 and 150 m (Fig. 8b), implying that saline Atlantic-origin Water has a pronounced influence at this depth.

The surface stratification starts developing in May and weakens from October onward (Figs. 8b and 15b). The sub-surface salinity maximum is only present from July to September (Fig. 8b). As the inflow is predominantly isopycnal (e.g., Rudels et al., 2002), the Atlantic-origin Water maximum appears as a convex bulge in TS space at inflow density (Fig. 8c). We used this characteristic signature in the late-summer profiles to identify the Atlantic-origin Water.
<figure><img src="image_10.png" /><figcaption>**Fig. 8.** Seasonal cycle of temperature and salinity in the upper 600 m from 1999 to 2019. Monthly mean profiles of temperature (a) and salinity (b) are shown by the colored lines. The shadings represent one standard deviation. The dashed black line at 50 m marks the depth of the layer most strongly impacted by seasonal variability. Panel (c) shows the monthly mean profiles in T S space.</figcaption></figure>

When convection begins, the Atlantic-origin Water maximum is eroded and becomes part of the mixed layer. In this way, the Atlantic-origin Water modifies the mixed-layer properties. Variability of the hydrographic properties of the Atlantic-origin Water will thus have an imprint on the mixed-layer development in the central Greenland Sea (e.g. Lauvset et al., 2018; Bashmachnikov et al., 2021). The long-term development of the Atlantic-origin Water maximum is consistent with the development of the mixed-layer properties (Fig. 9): temperature and salinity decreased until the mid-1980s and increased thereafter. The temperature and salinity increase of the Atlantic-origin Water maximum after the early 1990s first followed the $σ_θ = 28.02 \text{ kg/m}^3$ isopycnal until the early 2000s, when it became indistinguishable from those of the 1955–1975 period. In the following pentads, the density of the Atlantic-origin Water was reduced as the pronounced warming had a greater effect than the salinity increase. This particular change in the Atlantic-origin Water density after 2005 is reflected in the development of the mixed-layer density, which also started to decrease after 2005. We conclude that the development of the mixed-layer properties and the Atlantic-origin Water maximum strongly vary. Considering the small volume that constitutes the Atlantic-origin Water maximum (a layer shallower than 100 m, Fig. 8b, compared to mixed-layer depths of 500–1600 m, Fig. 5), it remains unclear to what extent the Atlantic-origin Water variability drives the mixed-layer development.

## Development of Greenland Sea Deep Water - the impact of advection from adjacent basins

The long-term development of the GSDW is inferred from the grid-ded hydrographic fields (Fig. 4). The freshening and cooling of the deepest mixed layers in the early 1980s were reflected in the development of the GSDW (Figs. 7 and 10). Since 1985, the GSDW has become warmer and saltier, following a nearly straight line through T S space (Fig. 10). The linear development indicates mixing with a warmer and saltier deep water mass. The candidates are NSDW and AODW, which are found in the adjacent deep basins and enter the Greenland Sea via troughs in the Mohn and Knipovich Ridges and through Fram Strait, respectively (Bönisch and Schlosser, 1995; Østerhus and Gammelsrød, 1999). The complex bathymetry and similarity of hydrographic properties make it challenging to understand the deep circulation. Geochemical tracer studies demonstrate that a continuous exchange between the deep water masses takes place (Smethie

et al., 1988; Swift and Koltermann, 1988; Bönisch and Schlosser, 1995; Bönisch et al., 1997; Langehaug and Falck, 2012), such that their properties equilibrate over time. Modification of hydrographic properties due to deep-water formation in the Greenland Sea and Eurasian Basin thus lead to successive adjustments of the deep water masses (Smethie et al., 1988; Heinze et al., 1990). The GSDW has converged toward AODW properties since the cessation of deep convection in the Greenland Sea (Bönisch et al., 1997; Meincke et al., 1997; Karstensen, 2005; Somavilla et al., 2013). A similar development occurred for the NSDW (Dickson et al., 1996), which matched the AODW properties around 2010 (Fig. 10). Mixing between GSDW and NSDW must also be considered a possible explanation for the hydrographic changes in the GSDW. A density-driven current from the Greenland Sea to the Norwegian Sea through the 2200 m-deep Jan Mayen Channel reversed its flow direction in the 1990s (Østerhus and Gammelsrød, 1999) due to the substantial density decrease in the Greenland Sea at 2000 m depth (Dickson et al., 1996). Volume transport estimates derived from chemical tracers and changes in hydrographic properties of the GSDW indicate that the flow of AODW into the deep Greenland Sea increased following the cessation of bottom-reaching convection (Somavilla et al., 2013). This may partly be attributed to additional inflow of AODW via the Norwegian Sea after the flow through the Jan Mayen Channel reversed. We cannot determine the relative contributions of NSDW and AODW to the development of the GSDW since 1985 as the hydrographic properties of both water masses are very similar (Fig. 10).

## Development of Greenland Sea Deep Water - the impact of geothermal heating

Geothermal heating has a non-negligible influence on both regional ocean bottom temperatures (e.g., Emile-Geay and Madec, 2009; Purkey and Johnson, 2010; Wunsch and Heimbach, 2014) and the large-scale abyssal circulation by weakening the abyssal stratification and increasing the circulation of deep water (e.g., Adcroft et al., 2001; de Lavergne et al., 2016; Downes et al., 2016). The trend of the mean global ocean temperature was estimated to $0.011 \pm 0.001 \frac{^\circ C}{\text{decade}}$ for the period 1994-2013 (Wunsch, 2018). This corresponds to a 20-year mean heating rate of $0.48 \pm 0.10 \frac{W}{m^2}$, of which nearly 20% arose from geothermal heating. To quantify the contribution from geothermal heating to the temperature increase of the GSDW, we calculated the mean heating inside the central Greenland Sea (thick black line in Fig. 3). We
<figure><img src="image_11.png" /><figcaption>**Fig. 9.** Development of the hydrographic properties of the Atlantic-origin Water maximum over the 70-year period. The colored circles represent the mean hydrographic properties of Atlantic-origin Water present in the central Greenland Sea during summer (July, August, September) averaged over the indicated periods (the black errorbars mark the standard deviations of the hydrographic properties). The periods match those used for determination of mixed-layer properties (Fig. 7). The colors show the average depth of the Atlantic-origin Water maximum. For comparison, the mixed-layer properties from Fig. 7 for the same periods are indicated by the blue dots. The light-gray isopycnals are referenced to surface pressure $\sigma_g$, the black 30.448 $\frac{\text{kg}}{\text{m}^3}$ isopycnal is referenced to 500 dbar and marks the separation of intermediate and deep water masses.</figcaption></figure>

<figure><img src="image_12.png" /><figcaption>**Fig. 10.** Development of the GSDW properties based on the gridded fields (Fig. 4) over the 70-year period. The colored circles indicate the shallowest depth of GSDW within each pentad. The letters “FS” represent AODW properties measured between 1900 m and 2100 m in western Fram Strait. The letters “NB” represent hydrographic properties of NSDW sampled near Ocean Weather Station M in the Norwegian Basin. Their developments are indicated by the dashed lines. All data points represent five-year averages except the first entries from western Fram Strait and Station M, which represent the average properties from 1950 to 1990. The boxes indicate the 10th–90th percentiles of GSDW properties for each pentad. The colors of the boxes indicate time and match the colors of the data points “NB” and “FS”. The standard deviations for “NB” and “FS” are not shown, but are of the same order of magnitude as the standard deviation of the GSDW. The isopycnals are referenced to 500 m, the $\sigma_{0.5} = 30.448 \frac{\text{kg}}{\text{m}^3}$ isopycnal separates intermediate and deep water masses.</figcaption></figure>

obtained an average of $111 \pm 83 \frac{\text{mW}}{\text{m}^2}$ with a maximum of $355 \frac{\text{mW}}{\text{m}^2}$, based on measurements distributed evenly across the central Greenland Sea. The age of the seafloor inside the central Greenland Sea ranges from $9 \cdot 10^6$ years to $56 \cdot 10^6$ years. For comparison, the global mean heat flux for oceanic crust in the same age span was $95 \frac{\text{mW}}{\text{m}^2}$, which is in good agreement with our estimate for the central Greenland Sea. Applying the mean basal warming for the central Greenland Sea over a 1500 m thick layer, we estimate a temperature increase of $0.005 \pm 0.004 \frac{^\circ\text{C}}{\text{decade}}$.

for the GSDW, which is an order of magnitude lower than the observed warming. Therefore, geothermal heating is not a leading-order cause for the GSDW temperature increase since the mid-1980s.

### Deep variability before the cessation of bottom-reaching convection

The $\sigma-\pi$ distance to 750 m, which is a measure of vertical homogeneity, showed most pronounced variability around 2500 m depth (Fig. 4d).
Years with a high degree of similarity within the water column cor-
respond to years when bottom-reaching convection occurred. In years
with high σ-π distance, hence low similarity, the hydrographic proper-
ties at depth changed independently of those closer to the surface. This
indicates modification due to enhanced lateral mixing, centered around
2500 m depth. After the cessation of bottom-reaching convection, the
development of the GSDW properties was dominated by mixing with
AODW and NSDW. The depth where the salinity of the GSDW first
started to increase in the 1990s corresponds approximately to the depth
of most pronounced variability in σ-π distance (Fig. 4d and b). We
may conclude that the variability of the σ-π distance around 2500 m
depth reflects an imprint of inflowing AODW and NSDW, also before
the cessation of bottom-reaching convection. Recurring incidences of
bottom-reaching convection eroded the signature of lateral inflow be-
fore 1985 (Fig. 4d and 5). Local overturning through bottom-reaching
convection must thus have taken place at sufficient volume to maintain
the cold, fresh signature of the GSDW before 1985.

Export and formation rate of Greenland Sea Deep Water

The density of the GSDW was greater than σ<sub>0.5</sub> = 30.46 kg/m<sup>3</sup> through-
out the entire 70-year period, which is substantially denser than the
isopycnal distinguishing deep and intermediate water masses (Fig. 10).
While the stratification maximum is included in the GSDW according to
this definition, the pronounced stratification rather marks the transition
between GSAIW and GSDW than indicating a homogeneous water mass
as is found below. Hence, the base of the stratification maximum is
a better measure of the depth range of the GSDW than the σ<sub>0.5</sub> =
30.448 kg/m<sup>3</sup> isopycnal after 1990 (Fig. 4e). During the 1990s, the volume
of GSDW gradually decreased as indicated by the descent of the strat-
ification maximum (Fig. 4c and e). The deepening of the stratification
maximum at an approximate rate of 1000 m/decade is consistent with a
net export of 0.6 Sv of GSDW. The constant rate of descent (Fig. 4e)
through the 1990s suggests that the GSDW export rate was relatively
stable, even after the cessation of bottom-reaching convection. If the
export rate of GSDW did not change when bottom-reaching convection
ceased (as suggested by Meincke et al., 1997), we may assume that
GSDW was formed at the same rate before 1985 to maintain a steady
reservoir. This estimate agrees well with a GSDW-formation rate of
about 0.5 Sv derived from tracer experiments (Smethie et al., 1988;
Bönisch and Schlosser, 1995). This estimate does not take into account
the warming-induced isopycnal descent, so should be considered an
upper limit.

To summarize, two modes of convection in the central Green-
land Sea can be distinguished. While the thermobaric effect facili-
tated bottom-reaching convection between the mid-1960s and the mid-
1980s, convection in the recent decades occurred through thermally-
driven gradual mixed-layer deepening. Between the two convective
states, low upper-ocean salinity led to the formation of a stratification
maximum which limited convection to the top 1000 m of the water
column. The recent temperature increase led to a substantial density
reduction despite a concurrent salinity increase. This resulted in a
change in stratification, from a transition-ocean to an α-ocean, and to
the formation of a new class of GSAIW. Since the formation of the
stratification maximum in the mid-1980s, the deep and intermediate
water masses of the Greenland Sea have been separated and developed
independently. Subsequent changes in the GSDW properties were pri-
marily driven by mixing with deep water masses from adjacent basins
while the recent temperature and salinity increase of the GSAIW closely
corresponds to the development of the Atlantic-origin Water that is
mixed into the central Greenland Sea. To understand the hydrographic
changes from a broader perspective, the development of the sea-ice
extent and atmospheric forcing over the 70-year period will be assessed
in the following.

<figure><img src="image_13.png" /></figure>

**Fig. 11.** Marginal ice zones of the three sea-ice patterns representative for the Greenland Sea. A self-organizing map algorithm was applied to daily February sea-ice concentration fields to identify three ice patterns. For each pattern, the average of all members resembles a typical sea-ice concentration field. The colored patches indicate the marginal ice zone (the area with 15–50% sea-ice concentration) for each pattern. The central Greenland Sea is outlined in orange. The 1000 and 3000 m isobaths are shown in light gray.

4. Atmosphere-ice-ocean interaction

Heat loss from the ocean to the atmosphere plays a critical role in water-mass transformation. Sea ice restricts heat exchange between the two fluids, but at the ice edge, where cold and dry air first meets the relatively warm ocean surface, turbulent heat fluxes are particularly strong (Marshall and Schott, 1999; Spensberger and Spengler, 2021). The location of the marginal ice zone is therefore of central importance to understand water-mass transformation in the Greenland Sea.

4.1. Ice-edge retreat

Over the 70-year period, the sea-ice extent in the Greenland Sea has varied substantially on both sub-seasonal and interannual time scales. The large-scale shape of the ice cover, which determines the length and width of the marginal ice zone, impacts the heat exchange between the ocean and the atmosphere. We applied a self-organizing map technique (Section 2.10) to daily sea-ice concentration fields from February, typically the month with most extensive sea-ice cover, to identify three distinct spatial patterns dominant in the Greenland Sea over the 70-year period (Fig. 11). The pattern we refer to as *open ocean* has a narrow marginal ice zone along the Greenland shelf break, leaving most of the central Greenland Sea ice free. *Odden* designates an ice tongue that reached far into the Greenland Sea. Its shape followed the Jan Mayen Current which branches off the EGC north of the West Jan Mayen Ridge, transporting relatively fresh and cold surface water into the central Greenland Sea (Shuchman et al., 1998; Wadhams, 1999). Local sea-ice formation facilitated by the presence of Polar Water at the surface and import of sea ice along the path of the Jan Mayen Current were both considered causes for the appearance of *Odden* (Shuchman et al., 1998; Wadhams and Cosimo, 1999). Accompanying *Odden* was an ice-free bay called Nordbukta to its north. Nordbukta, nearly surrounded by sea ice like a vast polynya, was subject to particularly strong heat fluxes that may have destabilized the water column (Germe et al., 2011; Vinje et al., 2002) and triggered deep convection (Visbeck et al., 1995; Marshall and Schott, 1999; Vinje et al., 2002). Brine rejection due to local sea-ice formation in Nordbukta was considered a plausible mechanism to initiate salinity-driven convection in the Greenland Sea (Rudels, 1986, 1990). The area classified as marginal ice zone when *Odden* was the dominant pattern was about five times greater than in winters characterized by the *open-ocean* pattern (Fig. A5.1). The width of the marginal ice zone was comparable in the *large ice extent* and *Odden* periods, but the shape of *Odden* caused a particularly wide
<figure><img src="image_14.png" /><figcaption>**Fig. 12.** Frequency of occurrence of the three sea-ice patterns shown in Fig. 11 in February over the 70-year period. The top panel shows the dominant pattern for each February. The color shading indicates prolonged periods characterized by each pattern.</figcaption></figure>

marginal ice zone in some parts of the central Greenland Sea (Figs. 11 and A5.1). The major difference between *Odden* and the *large ice extent* patterns is thus the absence of Nordbukta.

The annual distribution of the three sea-ice patterns (Fig. 12) demonstrates a clear trend of diminishing sea-ice extent in the central Greenland Sea, from almost complete ice cover to the nearly ice-free conditions characterizing the last two decades. In this context, the *Odden* pattern represents a transition state between high and low ice coverage. For only a few winters between the late 1970s and the late 1990s was this the dominant pattern (Fig. 12).

After the dominant pattern was determined for each year (Fig. 12), a running-mean filter with a width of 21 years was applied to the time series. Three distinct phases, each associated with one of the spatial sea-ice patterns, emerged. The resulting periods are 1950–1973, characterized by *large ice extent*; 1974–1998, when *Odden* was present; and 1999–2020, when the *open-ocean* pattern dominated. After satellite observations of sea-ice extent commenced in 1979, the spatial and temporal resolutions of the sea-ice data increased substantially. This resulted in a sharp transition in data quality, which impacts the sea-ice and heat-flux distributions from the ERA5 reanalysis. However, the transition from *large ice extent* to *Odden* in 1973/1974 does not coincide with the increase in data quality in 1979. This indicates that the patterns are representative of natural variability rather than changes in measurement technology. The transition from bottom-reaching to intermediate convection occurred during the *Odden* phase. The presence of *Odden* may have impacted convection in the Greenland Sea in many ways that will be examined further in the following sections.

## 4.2. Changes in heat loss

### Total heat loss

As a consequence of reduced sea-ice extent, the total heat loss from the increasingly ice-free central Greenland Sea to the atmosphere increased substantially over the 70-year period (Fig. 13). A substantial portion of heat loss takes place at the beginning of winter, in November and December. In some years, heat loss in these months constituted more than 50% of the total November to April heat loss (Fig. A6.1). We therefore consider the full six months from November to April when quantifying total winter heat loss. We only include turbulent heat fluxes in the analysis. Longwave radiation was not taken into account because both its contribution to the total heat loss and its interannual variability is low compared to the turbulent heat fluxes (e.g. Valdivieso et al., 2017).

While the ice-free portion of the central Greenland Sea increased approximately sixfold according to the transition from *large ice extent* to *open ocean*, the total heat loss only doubled over the 70-year period (Fig. 13). Several factors contribute to this discrepancy. Most (77%) of the variance in the heat exchange between the ocean and the atmosphere can be attributed to the variability in the ice-free area, which correlate at a 99% significance level. The interannual variability of the ice extent in the Greenland Sea was high, particularly before the Greenland Sea became predominantly ice-free (Fig. 13). This is a consequence of both the distribution of sea-ice extent within each pattern and the distribution of ice patterns within each winter. Furthermore, the heat loss is unevenly distributed across the ice-free area. The greatest heat loss takes place close to the ice edge (Fig. 14 and Spensberger and Spengler, 2021). While most of the ice-free area in the Greenland Sea was in the vicinity of the ice edge during the *large ice extent* phase, the currently ice-free areas are on average farther away from the ice edge. Consequently, the reduced sea-ice extent did not lead to a linear increase in heat loss with ice-free area. In the central Greenland Sea, the median winter heat loss from the ice-free area decreased by about $17 \pm 12 \text{ m}^2$, which corresponds to a 16% decrease over the 70-year period relative to 1950 according to a linear-regression analysis (Fig. S1.1). In the Greenland Sea Gyre (Fig. 3), located to the east of the greatest heat fluxes under *open-ocean* conditions (Fig. 14), the ocean-to-air heat flux decreased by about 20% between 1997 and 2014 (Moore et al., 2015). A main factor preventing the total heat loss from increasing at the same rate as the ice-free area was the westward shift of the band of highest heat fluxes that is tied to the location of the ice edge (Fig. 14).

### Extreme heat loss events

Most (60–80%) of the total wintertime heat loss takes place during cold-air outbreaks (Papritz and Spengler, 2017). As such, we also consider the high end of the heat-flux distribution when quantifying changes in surface forcing. The average number of days with heat fluxes exceeding the 95th percentile was $39 \pm 12$ per winter (November to April). Over the 70-year period, there was no significant trend in the number of such days, but the events became weaker (Fig. A6.2). These extreme events, however, affected increasingly large areas toward the end of the record as the sea-ice extent declined. This resulted in an overall increase of the total heat loss that can be attributed to extreme events (Fig. A6.2). Extreme heat flux events primarily occurred near the ice edge: 95% of all extreme events had their intensity maximum within a 2 km distance from the 15%-sea-ice contour, 99% within 6 km. The locations of the extreme heat fluxes also shifted westward along with the ice edge (Fig. 14).
<figure><img src="image_15.png" /><figcaption>**Fig. 13.** Total wintertime (November to April) heat loss from the ice-free portion of the Greenland Sea from 1950 to 2020. The heat loss obtained during each ice pattern is indicated by the colored bars. The sum of the patterns constitutes the total heat loss for each winter shown by the solid blue line. The dashed blue line is the corresponding linear regression line. The mean ice-free area for each winter is indicated by the red solid line, the red dashed line shows the corresponding linear regression line. The mean ice-free area (sea-ice concentration < 15%) is averaged over the same period as the total heat loss.</figcaption></figure>

### 4.3. Trends and changes in the seasonal variability of the surface layer

Next we examine the hydrographic properties of the surface layer (uppermost 50 m) of the central Greenland Sea, which along with the magnitude of the atmospheric forcing affects the mixed-layer properties at the end of winter. Surface conditions in the summer can provide insights about the preconditioning of the central Greenland Sea for convection in winter. The surface layer has a pronounced seasonal cycle typical for high-latitude oceans. The surface layer is freshest and warmest in summer. The maximum salinity and lowest temperature, resulting in the highest density, are reached in late winter (Fig. 15). Large deviations from the mean were typically non-symmetric. When, for example, the winter mean temperature was near the freezing point, variability occurred predominantly at higher temperatures. Salinity variability was more pronounced toward lower salinities.

The ice patterns define three periods with different surface conditions: 1950–1973 (large ice extent); 1974–1998 (Odden); and 1999–2020 (open ocean) (Fig. 12). The mean seasonal cycle for each period was reconstructed from all available measurements in the uppermost 50 m (Fig. 15). During the first period in particular, when only few winter measurements were available, the interannual variability may have been substantially greater than that represented by the standard deviation of the available data.

The surface temperature was higher by at least 0.5 °C throughout the winter months in the open-ocean period compared to the two proceeding periods. The mean winter surface temperature was near the freezing point between January and April in the two early periods. During the Odden period, both the summer surface temperature and salinity were particularly low, implying that less cooling was required to reach the freezing point. The statistically significant lower summer temperature and salinity during the Odden period may have been caused by a cold and fresh upper-ocean anomaly propagating northward along the Norwegian slope during the late 1970s (Holliday et al., 2008).

Despite the colder temperature and lower salinity during summer in the Odden period, the seasonal temperature and salinity cycles were remarkably similar in the both periods when sea ice was present in the Greenland Sea. In the open-ocean period, the surface-temperature increase in spring started approximately one month earlier. As a result, the winter season has been substantially shortened in the recent

decades. The more rapid warming in spring can be related to the lack of sea-ice cover in winter: no melting was necessary before solar radiation could heat the surface layer.

The seasonal cycle in salinity is strongly influenced by the seasonal movement of the Polar Front, which separates the Polar from the Arctic domains, along with brine rejection and freshwater input due to sea-ice formation and melt, respectively. During summer, the Polar Front extends far offshore and lets fresh PW into the central Greenland Sea (Våge et al., 2018, and Fig. 16). During winter, strong northerly winds push the Polar Front toward the Greenland slope, resulting in higher salinity in the central Greenland Sea (Spall et al., 2021). In the large ice extent and Odden periods, brine rejection due to local sea-ice formation may account for some of the salinity increase between January and March that did not happen during the open-ocean period (Fig. 15). When deepening mixed layers first tap into the Atlantic-origin Water, ventilation of warm and saline water increases the mixed-layer temperature and salinity. This may explain the large salinity increase in early winter (Fig. 15b). During the open-ocean period, the surface temperature was for the most part too high for local sea-ice formation to take place. Consequently, brine rejection did not impact the seasonal cycle. Melt of imported sea ice may have contributed to the salinity reduction in spring and early summer. As a result, the seasonal cycle in salinity was less pronounced in the open-ocean period than before.

The shortened cooling season may have a considerable impact on the product of water-mass transformation in the Greenland Sea. The onset of warming ends dense-water formation earlier in the season, such that we may expect lower surface density by the end of winter. To estimate the changes in density due to a shortened cooling season alone, the effect of changed salinity needs to be eliminated. For doing so, we constructed the seasonal density cycle based on the surface temperature from the open-ocean period and the mean surface salinity from the previous periods. The resulting density cycle has its maximum in March. In May, the surface density was reduced by up to 0.09 $kg/m^3$ compared to the previous periods due to the earlier onset of warming. The shorter convective season combined with the substantially higher surface temperatures of the recent decades may, in the extreme case, have caused a temporary shut-down of convection, but the temperature increase was compensated by a high surface salinity, resulting in surface densities comparable to the early periods. The salinity increase led
<figure><img src="image_16.png" /><figcaption>**Fig. 14.** Spatial distribution of the largest heat fluxes and locations of maximum heat loss for each ice pattern in February. Left panels (a, c, e): composite means of the 5% greatest heat fluxes for each February that constituted the sea-ice pattern. Right panels (b, d, f): corresponding spatial distributions of the 5% greatest heat fluxes. For each day the location of maximum heat flux was calculated. The color indicates the relative frequency of each location with heat fluxes higher than the 95th percentile (412 W/m²). For each pattern, N is the number of days in February with heat fluxes greater than the 95th percentile. The orange lines indicate the 15% and 50% sea-ice contours. The dark contour outlines the central Greenland Sea.</figcaption></figure>

to a density increase of approximately 0.15 $\frac{\text{kg}}{\text{m}^3}$. In the Greenland Sea, deep and intermediate convection only took place when the salinity was sufficiently high (Fig. 6). As such, salinity contributed to set the convective state of the Greenland Sea over the 70-year period. We conclude that the shallow-convection phase in the late 1980s and early 1990s was probably not caused by changes in the seasonal cycle, but rather by a low sea-surface salinity that inhibited convection.

### The seasonal movement of the Polar Front

The Polar Front has been identified by the 34.66 $\frac{\text{g}}{\text{kg}}$ salinity contour (Swift and Aagaard, 1981; Brakstad et al., 2023). It cuts across the central Greenland Sea in summer, while during winter it is located much farther to the west, close to the east Greenland shelf (Fig. 16). Over most of the central Greenland Sea, the surface salinity increases in winter compared to late summer. The most pronounced seasonal salinity increase is caused by this lateral migration of the Polar Front. The median change to the west of the Polar Front is greater than 0.6 $\frac{\text{g}}{\text{kg}}$ in both the *Odden* and *open-ocean* periods (Fig. 16a and b). To the east of the Polar Front, the median salinity change is comparatively subdued, on the order of 0.1 $\frac{\text{g}}{\text{kg}}$, and may to some extent be caused by ventilation of Atlantic-origin Water. When the mixed layer first taps into the Atlantic-origin Water, the salinity is uniformly distributed across

the entire mixed layer and thus increases the surface salinity (Ronski and Budéus, 2005). If we assume that ventilation occurs uniformly across the central Greenland Sea, a salinity change of approximately 0.5 $\frac{\text{g}}{\text{kg}}$ can be attributed to the migration of the Polar Front, while the remaining 0.1 $\frac{\text{g}}{\text{kg}}$ to close the budget may stem from the ventilation of Atlantic-origin Water.

The easternmost extent of the Polar Front was reduced from the *Odden* to the *open-ocean* period (Fig. 16). This change is consistent with the reduction of winter sea-ice extent and coincides with the transition from bottom-reaching to intermediate convection. As such, it is another expression of the systematic change that took place in the Greenland Sea in the late 1990s and early 2000s.

To summarize, the sea-ice extent in the Greenland Sea was classified into three patterns. Over the 70-year period, a change from *large ice extent* to *open ocean* took place, with *Odden* as a transient phase in between, from 1974 to 1998. The phase of bottom-reaching convection began under *large ice extent* conditions, and ceased during the *Odden* period. In both of these periods, the surface temperature was particularly low in late winter, which is a necessary condition for thermobaric convection. Initially, the low surface salinity during the *Odden* phase did not prevent bottom-reaching convection. Instead, the low-salinity surface water was dispersed across the mixed layer and
<figure><img src="image_17.png" /><figcaption>**Fig. 15.** Seasonal cycle of (a) temperature, (b) salinity, and (c) density in the upper 50m of the central Greenland Sea for the three periods defined by the ice patterns in Fig. 12: 1950–1973 (blue, large ice extent), 1974–1998 (pink, Odden), and 1999–2020 (orange, open ocean). The solid lines show the seasonal cycle as monthly means of all available measurements in each period. The standard deviation indicated by the shaded areas was calculated separately for measurements above and below the mean to account for asymmetric variability.</figcaption></figure>

<figure><img src="image_18.png" /><figcaption>**Fig. 16.** Sea surface salinity change from August-October (ASO) to February-April (FMA) for the Odden (1974–1998) and open-ocean (1999–2019) periods. Panels (a) and (b) map the summer-to-winter salinity change. Positive values indicate that salinity in FMA is higher than in ASO. The gray contour indicates the location of the Polar Front, defined as the 34.66 ± 0.15 kg/m³ salinity contour (Swift and Aagaard, 1981; Brakstad, 2023a), in ASO. Panels (c) and (d) show histograms of sea surface salinities for ASO (red) and FMA (blue) for the periods mapped in the panels above. The stars mark the median salinity for each season. Only salinity measurements from the upper 20m were included.</figcaption></figure>

led to a freshening of the GSDW. In the second half of the *Odden* period, further reductions in salinity limited convection and led to the formation of the stratification maximum. Toward the end of the *Odden* period, the sea-ice cover in the central Greenland Sea had nearly vanished. The resulting increase in total heat loss was, however, not sufficient to compensate for the temperature increase. High surface

temperatures in the *open-ocean* period may have prevented the presence of sea-ice during winter, which in turn led to an earlier onset of warming in spring. Our results thus emphasize the coupling between the ocean, cryosphere, and atmosphere, and illustrate how the 70-year development of the Greenland Sea hydrography cannot be considered in isolation.
# 5. Discussion

## Bottom-reaching convection in the Greenland Sea

We identified several profiles showing bottom-reaching convection between 1965 and 1985. Furthermore, this period was characterized by exceptionally weak stratification and a homogeneous water column compared to the rest of the 70-year period (Fig. 4d and e). We conclude that no bottom-reaching convection took place between 1950 and the mid-1960s or after the mid-1980s. Helland-Hansen and Nansen (1909) inferred bottom-reaching convection from observations of exceptionally weak stratification in the Greenland Sea in the early 1900s. We surmise that the Greenland Sea has previously oscillated between states of increased and reduced GSDW formation. We found that thermobaric convection may have facilitated bottom-reaching convection between the late 1960s and early 1980s (Fig. 6). Thermobaric convection can only take place when mixed layers are sufficiently cold. Our results therefore support the hypothesis that linked bottom-reaching convection in the 1960s to a period of exceptionally cold winters, which resulted in particularly cold mixed-layers (Aagaard, 1968; Alekseev et al., 2001). This also supports the notion that the cooling of the GSDW in the mid-1960s was a sign of GSDW ventilation due to bottom-reaching convection (Meincke et al., 1992).

While we find evidence for bottom-reaching convection until the mid-1980s, reduced levels of dissolved oxygen suggest that substantial renewal of GSDW has not occurred since the end of the 1970s/beginning of the 1980s (e.g., Schlosser et al., 1991; Rhein, 1991; Bönisch and Schlosser, 1995). Others report evidence for bottom-reaching convection as late as winter 1989 (Alekseev et al., 1994; Meincke et al., 1997) and postulate that bottom-reaching convection ceased in 1990. Taking into consideration the mixing with a warmer and more saline deep water mass that occurs around 2500 m depth (Figs. 4b and d; Section 3.1), the sustained low salinity during the late 1980s and the relatively low $\sigma-\pi$ distance between 750 m depth and the bottom support the hypothesis that bottom-reaching convection took place until 1990. The weak but notable increase of $\sigma-\pi$ distance in the late 1980s compared to the early 1980s, however, may indicate that the GSDW was at most partially renewed. We date the cessation of substantial GSDW production to the mid-1980s, when local overturning was still vigorous in the central Greenland Sea. After 1985, the development of the GSDW properties was primarily driven by isopycnal mixing with NSDW and AODW (Fig. 10).

## Hydrographic preconditioning for bottom-reaching and thermobaric convection

Thermal and haline convection were considered possible processes for dense-water formation in the Greenland Sea (e.g., Aagaard and Carmack, 1989; Meincke et al., 1997). During thermal convection, the mixed layer deepens gradually due to surface cooling. Haline convection is a multi-step process in which first sea-ice formation at the top of a fresh and cold surface layer leads to the formation of high-salinity plumes. When the plumes reach a warmer and more saline subsurface layer, ventilation brings the warmer water to the surface and the newly formed sea ice melts (Rudels, 1986, 1990). In a transition-ocean, both temperature and salinity contribute to the stratification. We can infer that neither thermal nor haline convection strictly dominated overturning to mid-depth in the Greenland Sea before the formation of the stratification maximum. For the recent decades, due to the transition into an $\alpha$-ocean (Fig. 4f), we can expect thermal convection to dominate mixed-layer deepening. Note that even in a temperature-dominated regime, sufficient salinity is needed to achieve intermediate convection (Fig. 6). For the Greenland Sea, convection exceeding 1000 m took place at a minimum mixed-layer salinity of 35.03 $g/kg$ (Fig. 6b). Brakstad et al. (2019) found that a near-surface salinity greater than 34.88 $g/kg$ the preceding summer was necessary for the formation of GSAIW. The difference between the summer value of 34.88 $g/kg$ and the winter value of 35.03 $g/kg$ may be achieved by admixture of salt from the

layer of Atlantic-origin Water during convection (Ronski and Budéus, 2005) which is on the order of $0.1 g/kg$ (Fig. 16 and Section 4.3). Below 2000 m, where the Greenland Sea has been temperature-stratified throughout the 70-year period, convection was thermally driven. This is consistent with the narrow band of particularly low mixed-layer temperatures for bottom-reaching convection compared to the small but depth-independent range of mixed-layer salinity. Thermobaric convection is a special case of thermal convection that occurs in cold, weakly-stratified oceans once the thermobaric depth has been reached. The relatively shallow (approximately 1500 m deep) thermobaric depth before 2000 facilitated the occurrence of thermobaric convection in the Greenland Sea between 1965 and 1985, when particularly cold mixed-layers formed.

## The "halocline catastrophe"

Increased amounts of near-surface freshwater can lead to the development of a halocline which prevents mixing between the upper layer and the deeper ocean. If a halocline develops in a convection region, the subsequent shut-down of deep convection is referred to as a "halocline catastrophe" (Rooth, 1982). The decoupling of the GSDW from the upper ocean took place due to low upper-layer salinities that prevented deep convection (Fig. 7 and Brakstad et al., 2019). This is emphasized by the salinity stratification dominating the temperature stratification when the stratification maximum formed in the 1990s (Fig. 4e). The separation of the Greenland Sea into upper and lower strata due to this salinity stratification had further consequences on the vertical distribution of heat within the water column. Two main heat sources for the central Greenland Sea are mixing with Atlantic-origin Water near the surface (Svingen et al., 2023) and with deeper water masses (AODW and NSDW) near the bottom (Fig. 10). As the Atlantic-origin Water is substantially warmer than the two deep water masses, the majority of the heat is added to the upper part of the Greenland Sea. When bottom-reaching convection took place, heat imported into the Greenland Sea was distributed across the entire water column. After the stratification maximum formed, heat added to the upper ocean remained within the comparatively thin upper layer that was ventilated during winter convection. Consequently, the same amount of heating distributed over a smaller volume led to an upper-ocean temperature increase. An additional warming of the Atlantic-origin Water further magnified this effect (Fig. 9). The temperature increase of the layer below the stratification maximum was subdued compared to the upper layer (Fig. 7). This led to an increasing temperature difference between GSAIW and GSDW. When the mixed-layer salinity increased in the 2000s, the salinity stratification was weakened, but the temperature difference was sufficiently strong to maintain the stratification maximum between the intermediate and deep water masses (Fig. 4f and e). The presence of the stratification maximum thus had a self-perpetuating effect by preventing a vertically homogeneous temperature increase. We conclude that the freshening of the 1980s led to an interruption of bottom-reaching convection, which resulted in a stratification-preserving feedback mechanism. The cessation of deep convection in the Greenland Sea may therefore be classified as a "halocline catastrophe" (Aagaard and Carmack, 1989). The large temperature increase of the recent decades has since transformed the Greenland Sea into a source of intermediate rather than deep water masses.

## Atlantification of the Greenland Sea?

Aspects of the development in the central Greenland Sea, such as the recent temperature and salinity increases, may be interpreted as signs of "Atlantification". As for the Barents Sea (Årthun et al., 2012) and parts of the Arctic Ocean (Polyakov et al., 2017), Atlantification can be traced to the inflow of warmer and more saline Atlantic-origin Water and resulted in a reduction of the winter sea-ice extent. The regime-shift in stratification in the Greenland Sea from a transition- to an $\alpha$-ocean, typical for the Atlantic domain, may justify application of the
term Atlantification. While the Atlantification of the Arctic Ocean led to a weakening of the stratification and thus facilitated deeper mixed layers, the temperature and salinity increase in the central Greenland Sea had the opposite effect and contributed to the strengthening of the stratification maximum, limiting convection to the upper 2000 m.

## Great Salinity Anomalies

Approximately once per decade since the 1970s, Great Salinity Anomalies propagated around the Subpolar Gyre of the North Atlantic (Dickson et al., 1988; Belkin et al., 1998; Belkin, 2004; Holliday et al., 2020). Eventually, the anomalies crossed the Greenland-Scotland Ridge and followed the Norwegian slope toward Fram Strait (Dickson et al., 1988; Belkin et al., 1998; Belkin, 2004; Holliday et al., 2008). Although the fresh anomalies typically weaken as they progress northward, the signals are still detectable in Fram Strait (Holliday et al., 2008). As the Greenland Sea mixed-layer salinity is strongly influenced by the salinity of the Atlantic-origin Water (Fig. 9), the Great Salinity Anomalies likely reduce salinities in the Greenland Sea. Following Dickson et al. (1988), the freshening of the 1980s may thus be interpreted as the arrival of the 1970s' Great Salinity Anomaly in the central Greenland Sea. The Great Salinity Anomalies of the 1970s, 1980s, and 1990s were first detected either in the Iceland Sea (Dickson et al., 1988) or at the southern tip of Greenland Belkin et al. (1998), Belkin (2004). While their amplitudes were similar in the West Greenland Current, the anomalies of the 1980s and 1990s were substantially weaker in the Nordic Seas (Holliday et al., 2008). This may help explain why the more recent Great Salinity Anomalies did not leave an equally clear imprint on the Greenland Sea. It is likely that some of the Great Salinity Anomalies originated in the Arctic Ocean, which may be considered a vast freshwater reservoir (e.g., Giles et al., 2012). A large portion of southward-flowing freshwater from the Arctic Ocean into the Nordic Seas is exported on the Greenland shelf through Fram Strait (Karpouzoglou et al., 2023). Most of the fresh water on the Greenland shelf is trapped in the East Greenland Coastal Current until it reaches the southern tip of Greenland Foukal et al. (2020), while only smaller amounts are diverted into the deep basins along its southward pathway (Duyck and de Jong, 2023). Freshwater anomalies originating in the Arctic Ocean may thus propagate southward on the Greenland shelf without noticeably impacting the central Greenland Sea.

## Large-scale impacts of the cessation of bottom-reaching convection

The cessation of bottom-reaching convection in the Greenland Sea led to a reorganization of the exchange of deep water masses between the Greenland Sea, Norwegian Sea, and Arctic Ocean. During the phase of active production, GSDW was exported into the other deep basins, such that the Greenland Sea acted as a source of deep water for the Arctic Mediterranean. In the 1990s, the supply of GSDW to the Arctic Ocean through Fram Strait stopped (Langehaug and Falck, 2012) and the flow through the Jan Mayen Channel reversed (Østerhus and Gammelsrød, 1999). This indicates the importance of water-mass transformation in the Greenland Sea for the deep circulation in the Arctic Mediterranean. When GSAIW started forming instead of GSDW, the export pathways from the Greenland Sea also adjusted. Since 2005, GSAIW has contributed substantially to the densest component of the overflows in Denmark Strait and Faroe Bank Channel (Huang et al., 2020; Brakstad et al., 2023), which supply the lower limb of the AMOC (e.g., Swift et al., 1980; Chafik and Rossby, 2019). The cessation of GSDW formation and subsequent production of an intermediate water mass thus led to a redistribution of water masses supplying the lower limb of the AMOC.

# 6. Conclusions

We reconstructed the hydrography of the central Greenland Sea over the 70-year period from 1950 to 2020 using observations from bottle measurements, CTDs, and Argo floats. After a phase of active bottom-water renewal between the mid-1960s and the mid-1980s, a stratification maximum formed and isolated the deep from the intermediate Greenland Sea. As a consequence, a new class of GSAIW formed (Brakstad et al., 2019) and has developed independently of the GSDW since the 1980s (Figs. 7 and 10) and the stratification changed from a transition- to an α-ocean (Fig. 4f). Coinciding with this change was a substantial reduction of the winter sea-ice extent, which resulted in an increase of the total heat loss from the Greenland Sea.

## Formation of the stratification maximum

The formation of the stratification maximum that isolated the GSDW from the upper water column (Fig. 4e and Karstensen, 2005; Latarius and Quadfasel, 2010; Brakstad et al., 2019) along with the formation of a new class of GSAIW (Brakstad et al., 2019) beginning in the 1990s was unique within the recent 70-year period in the Greenland Sea. This transition is crucial to our understanding of deep-water formation in general and of changes in the Greenland Sea convection in particular. A freshwater anomaly limited convection to the upper 1000 m in the late 1980s and early 1990s and resulted in the formation of the stratification maximum (Fig. 7 and Brakstad et al., 2019). As a consequence, the deep and intermediate water masses in the Greenland Sea developed independently toward higher temperatures and salinities. The more rapid warming of the GSAIW resulted in a faster density decrease, which enhanced the stratification maximum. The formation of the stratification maximum marks the transition from thermobaric to thermal convection in the central Greenland Sea.

## Short-term vs. long-term variability in dense-water formation

For the Greenland Sea Gyre, the heat loss has diminished by 20% since 1979 (Moore et al., 2015). For the entire central Greenland Sea, however, the total heat loss increased over the 70-year period due to an increase in the ice-free area (Fig. 13). The warming of the GSAIW can thus not be explained by reduced cooling. We suggest that the import of warmer Atlantic-origin Water was an important cause of the warming and density reduction of the GSAIW (Fig. 9). Annually, the total ocean-to-atmosphere heat loss through winter may determine the volume of water that can be transformed into a higher density class. The resulting water mass product, however, depends on whether haline or thermal convection dominates, which in turn is determined by the initial early-winter stratification. The long-term variability of water-mass transformation is therefore regulated by the hydrographic structure of the Greenland Sea.

## The effect of Odden/Nordbukta on mixed-layer deepening

Though the transition between bottom-reaching and intermediate convection occurred during the Odden phase, no causal link between the formation of Odden/Nordbukta and the occurrence of deep convection was uncovered. The formation of Odden was probably facilitated by the low surface salinity characterizing this period (Fig. 15), which eventually caused the cessation of bottom-reaching convection. Despite the opening of Nordbukta, which could have increased the potential for convection, no noticeable polynya effect was found. In polynyas, wind-driven sea-ice export repeatedly exposes the ocean surface to the atmosphere, such that continued heat loss and freezing is facilitated (e.g., Visbeck et al., 1995). However, neither did brine release due to sea-ice formation after sea-ice export from Nordbukta increase the salinity noticeably, nor was the effective heat flux during extreme events greater than for the other two phases (Fig. A6.2). As such, the presence of Odden/Nordbukta was not a requirement for the occurrence of bottom-reaching convection.
## Outlook

A further increase in the northward heat transport into the Barents Sea and Arctic Ocean is expected over the next decades (Årthun et al., 2019; Shu et al., 2022). Over the past decade, a warming and freshening of the Atlantic Water in the Norwegian Sea was documented (Mork et al., 2019). If this signal propagates into the Greenland Sea, combined with further increases in northward heat transport, reductions in the formation rate and density of GSAIW may be expected. As the Greenland Sea is currently a main source of the densest portion of the overflow water through Denmark Strait (Huang et al., 2020; Brakstad et al., 2023), prolonged reductions in the Greenland Sea dense-water production may have important implications for the supply of dense water to the lower limb of the AMOC.

## CRediT authorship contribution statement

**Anna-Marie Strehl:** Writing – review & editing, Writing – original draft, Visualization, Methodology, Investigation, Formal analysis, Data curation, Conceptualization. **Kjetil Våge:** Writing – review & editing, Validation, Supervision, Methodology, Investigation, Conceptualization. **Lars H. Smedsrud:** Writing – review & editing, Supervision, Conceptualization. **Thibaut Barreyre:** Writing – review & editing, Resources, Investigation, Data curation.

## Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

## Data availability

Most of the data is already published and referenced in the manuscript (Table A1.1). The remaining unpublished cruise data will be made available upon request. The gridded fields showing the development of the Greenland Sea hydrography (Fig. 4) will be published at the Norwegian Marine Data Centre and can additionally be requested from the corresponding author.

## Acknowledgments

K. Våge was funded by the European Union (ERC, ROVER, 101088452).

## Appendix A

### A1. List of data sources

An overview over the hydrographic data sets included in the analysis is provided in Table A1.1.

### A2. Filtering and gridding of the hydrographic data

#### Initial quality control

As a first step, profiles showing density inversions greater than 0.01 kg/m³ were removed. The threshold was twice that of Våge et al. (2018) to account for the larger measurement uncertainties early in the record. In the profiles obtained before 2000, only measurements causing an inversion were deleted instead of the entire profile. Furthermore, a density-difference criterion in density-depth space was applied: Measurements deviating by more than 0.005 kg/m³ from all other data points at the same depth within the 70-year record were classified as outliers and removed from the data set. To account for density-compensating spikes in temperature and salinity, this maximum difference criterion was also applied in temperature-depth and salinity-depth spaces with

<table><caption>**Table A1.1**
List of included hydrographic data sets, their authors and doi link.</caption><thead><tr><th>Name of and/or region covered by data set</th><th>Reference</th></tr></thead><tbody><tr><td>Hydrographic and geochemical observations in the Nordic Seas between 1950 and 2019 ("Historical data set")</td><td>Brakstad et al., 2023<br>https://doi.org/10.21335/NMDC-1271328906</td></tr><tr><td>Climatological Atlas of the Nordic Seas and northern North Atlantic</td><td>Korablev et al., 2014<br>https://doi.org/10.7289/V5K64G16</td></tr><tr><td>Meteor cruise M8/1 (1988), Greenland Sea and Barents Sea Opening</td><td>Quadfasel, 2018<br>https://doi.org/pangaea.de/10.1594/PANGAEA.890188</td></tr><tr><td>POLARSTERN cruise ARK-XV/3 (1999), Greenland Sea, Fram Strait, and Arctic Ocean</td><td>Schauer, 2010<br>https://doi.org/pangaea.de/10.1594/PANGAEA.742657</td></tr><tr><td>Jan Mayen Cruise 2001, Greenland and Iceland Seas</td><td>Holfort, 2001<br>The data have not been published as a data set yet, but were described in Wadhams et al. (2002). We obtained the data from personal communication with Jürgen Holfort. The cruise report is available here:<br>https://www.ifm.uni-hamburg.de/workareas/experimental/berichte/jan_mayen.pdf</td></tr><tr><td>Lance Cruise 21-2001, Greenland Sea</td><td>Logemann, 2018<br>https://doi.org/pangaea.de/10.1594/PANGAEA.890880</td></tr><tr><td>Moorings J008-J034 and JP31/JP34 in the Greenland Sea, 1999–2009</td><td>Budéus et al., 2020<br>https://doi.org/pangaea.de/10.1594/PANGAEA.911001</td></tr><tr><td>NPI cruise FS2014, Fram Strait</td><td>Dodd, Divine, Granskog, Johansson, Kern, Pavlov, Peralta Ferriz, Stedmon and de Steur, 2022<br>https://doi.org/10.21334/npolar.2022.493ea7ad</td></tr><tr><td>NPI cruise FS2015, Fram Strait</td><td>Dodd et al., 2022b<br>https://doi.org/10.21334/npolar.2022.52ecdc98</td></tr><tr><td>NPI cruise FS2016, Fram Strait</td><td>Dodd et al., 2022c<br>https://doi.org/10.21334/npolar.2022.29c6e2c7</td></tr><tr><td>NPI cruise FS2017, Fram Strait</td><td>Dodd et al., 2022a<br>https://doi.org/10.21334/npolar.2022.44db5c55</td></tr><tr><td>NPI cruise FS2018, Fram Strait</td><td>Dodd, Chierici, Debyser, Fransson, Granskog, Hamar, Kern, Rhode Kier, Lautkötter, Stedmon, de Steur and Wefing, 2022<br>https://doi.org/10.21334/npolar.2022.2c646c2e</td></tr><tr><td>NPI cruise FS2019, Fram Strait</td><td>Dodd et al., 2022e<br>https://doi.org/10.21334/npolar.2022.5066a075</td></tr><tr><td>NPI cruise FS2020, Fram Strait</td><td>Dodd et al., 2022d<br>https://doi.org/10.21334/npolar.2022.5df344c6</td></tr></tbody></table>

thresholds of 0.01 °C and 0.005 kg/kg, respectively. With this method, spikes could be removed that caused density inversions of less than 0.01 kg/m³. After removing spikes in temperature, salinity, and density, further outliers were identified by applying a simple statistical filter in density-depth space. First, we calculated a median density profile using all measurements from the central Greenland Sea. Data points that deviated by more than 15 standard deviations from the median density profile at any depth were then identified and the corresponding profiles removed. The wide range was chosen to ensure removal only of unequivocal outliers, smaller deviations were within the bounds of long-term variability. Out of more than 10 500 profiles, only three profiles were removed by this filter.

#### Gridding

To facilitate a quantitative analysis, all of the temperature, salinity, and density profiles from the central Greenland Sea were interpolated onto a regular pressure-time grid using a Laplacian spline
<figure><img src="image_19.png" /><figcaption>**Fig. A2.1.** Temporal data availability and vertical resolution for the entire 70-year period. The spacing on the time axis reflects the temporal resolution applied when gridding the different periods. Left y-axis: number of profiles per time interval. Right y-axis: width of the Hanning-window used for smoothing the gridded fields after the interpolation. The mean vertical resolution of the profiles is given for the periods separated by the black lines: 1950–1972, 1973–2000, and 2001–2020.</figcaption></figure>

algorithm (Pickart and Smethie, 1998). The resolution of the grid was adjusted according to the amount of available data. While the sparse data in the beginning of the 70-year period only resolves interannual variability, the much larger number of profiles in the last 20 years is sufficient to resolve sub-seasonal signals. We therefore define three periods based on the amount of available data per time interval: 1950–1972, 1973–2000, and 2001–2020. For each of these periods, a temporal resolution was chosen such that the amount of data per grid point in time is similar for all three periods (Fig. A2.1). The resulting periods and their respective resolutions are listed in Table A.1.

To achieve coherent hydrographic fields representing the entire 70-year period, the three periods were gridded successively. As separate gridding of each period would introduce discontinuities at the junctions between periods, we applied an iterative process in which the resolution of the grid was increased in a step-wise manner. A visualization of the process is shown in Fig. A2.2. First, all available data were gridded on the lowest-resolution grid (6 months in time, 50 m in depth, see Table A.1). Then, the temporal resolution was increased to 2 months. To ensure sufficient data coverage of the first period while increasing the resolution, the gridded field from the first step was used as input instead of the actual data for this period. Finally, the temporal resolution was increased to 8 days and the vertical resolution doubled. Now only data from the last period was used, while the previously gridded field served as input data for the first two periods. The final resolution of all gridded fields used for the analysis is thus 8 days in time and 25 m in depth, but only the most recent period truly resolves variability at this frequency.

## Final quality control

To identify and remove high-frequency variability, resulting from the sampling of eddies or fronts that were not representative of the current conditions in the central Greenland Sea, we smoothed the gridded density field using a Hanning-window filter with a constant width of 4 years. Profiles causing differences between the smoothed and original density fields exceeding 2.5 times the density range at any

<caption>**Table A.1**</caption>

Periods and resolution used in the iterative gridding procedure. The parameters were chosen such that within each period the amount of data per grid point is similar (Fig. A2.1).

<table><thead><tr><td>Period</td><td>Vertical spacing</td><td>Temporal spacing</td></tr></thead><tbody><tr><td>1950–1972</td><td>50 m</td><td>6 months</td></tr><tr><td>1973–2000</td><td>50 m</td><td>2 months</td></tr><tr><td>2001–2020</td><td>25 m</td><td>8 days</td></tr></tbody></table>

depth level were discarded. Less than 2% of the profiles were removed by this procedure.

Spline interpolation may amplify minor deviations, particularly in regions with sparse data and near the boundaries. Some profiles from August 1962 and June 1983 contained a minimal instability at the bottom, which was intensified by the interpolation. The data points causing this instability were discarded. In total, ten data points were removed by this process.

## Smoothing

After the filtering was complete, the data were gridded again and smoothed using a Hanning-window filter to remove temporal variability on time scales beyond the resolution of the available measurements in the early periods. The width of the filter ranged from five years in the first period to less than one year in the final period (Fig. A2.1) to allow for seasonal variability. The width of the Hanning-window was constructed as a combination of two Gompertz functions (Gompertz, 1833). A Gompertz function describes a transition between two asymptotically constant states 0 and 1. The properties of the transition are determined by parameters $\beta$ (position of inflection) and $\gamma$ (slope of inflection), while a parameter $\alpha$ can be used to stretch the distance between the asymptotic states of the function before and after the inflection. The Gompertz function for the width of the Hanning-window filter G as shown in Fig. A2.1 is described by the following equation:

$$G(t) = 5 \cdot \left( 1 - \alpha_1 \cdot e^{-e^{-\gamma_1(t-\beta_1)}} - \alpha_2 \cdot e^{-e^{-\gamma_2(t-\beta_2)}} \right), \quad (1)$$
<figure><img src="image_20.png" /><figcaption>**Fig. A2.2.** Visualization of the step-wise gridding process. The step-wise increase in resolution is illustrated by plotting every 12th grid line in both time and depth. The red lines mark the transitions between the periods defined by different data availabilities as shown in Fig. A2.1. Above each panel, the input data for the respective periods are printed, periods for which the previously gridded field was used as input are highlighted in light blue.</figcaption></figure>

where $\alpha_1 = \frac{1}{5}$, $\alpha_2 = 1 - \alpha_1 - 0.125$, $\beta_1 = 1973 - 1950$, $\beta_2 = 1997 - 1950$, $\gamma_1 = 0.005$, $\gamma_2 = 0.009$, *t* represents time in years, and *e* is Euler's number.

## Error estimates

To assess the uncertainty of the gridded fields, 1000 samples of the data set were created following a bootstrap procedure. Each sample was gridded and the standard deviations of the complete set of samples were calculated for each field. The resulting standard deviations of the temperature, salinity, and density fields are shown in Figs. A2.3–A2.5. Apart from the upper 500 m, where a pronounced seasonal cycle adds substantial variability, the standard deviations are less than or equal to the measurement errors in the respective periods. The long-term variability is consistent with the original gridded fields, in the mean as well as in the minimum and maximum of the subsampled data. From this procedure, we conclude that the error and statistical variability in the gridded fields do not exceed the measurement error.

The standard deviations of temperature and density were enhanced around 2000 m between 2010 and 2020 (Figs. A2.5 and A2.3) and thus collocated with the stratification maximum (Fig. 4d). Since stratification is calculated from the vertical density gradient, this increased standard deviation may have amplified the stratification maximum. Argo floats only sample the upper 2000 m of the water column. Just below their maximum sampling depth, a consistently high stratification is indeed present during that period (Fig. 4d). This particular feature may not be real. However, the enhanced stratification both above and below 2000 m makes us confident that the stratification maximum is present during that period and that it is localized at approximately this depth. The continued independent developments of GSAIW and GSDW after 2010 (Figs. 7 and 10) further corroborate that the separation of the water masses caused by the stratification maximum persisted beyond 2010.
<figure><img src="image_21.png" /><figcaption>**Fig. A2.3.** Standard deviation of potential density $\sigma_3$ derived from gridding with 1000 bootstrap samples. The color scale is linear until the 95th percentile of standard deviations. The highest value indicates the maximum standard deviation, the second highest indicates the 98th percentile. As such, the standard deviation of density is never larger than 0.26 $kg/m^3$, but in 98% of the grid points the standard deviation is even smaller than 0.0114 $kg/m^3$ and 95% are smaller than 0.0045 $kg/m^3$. As in Fig. 4, each black line at the top indicates the presence of a profile, red and white colors indicate summer (May-Oct) and winter (Nov-Apr) months, respectively. At the bottom, profiles extending deeper than 2000 m are marked.</figcaption></figure>

<figure><img src="image_22.png" /><figcaption>**Fig. A2.4.** Standard deviation of Absolute Salinity derived from gridding with 1000 bootstrap samples. The color scale corresponds to that of density (Fig. A2.3).</figcaption></figure>

Uncertainty of measurements

The uncertainty of measurements improved substantially over the 70-year period. In Table A.2, an overview of characteristic uncertainties for different periods is provided.

A3. Differences between the EOS-80 and the TEOS-10 standards

The density of a water parcel differs slightly between the EOS-80 and the TEOS-10 standards. Due to the weak stratification in the Greenland Sea, even discrepancies at the 3rd decimal level can translate into a difference in depth of several tens of meters. As GSDW is defined by density, it is necessary to adjust the definition by Rudels et al. (2002), based on the EOS-80 standard, to the TEOS-10 standard used here. The differences vary with temperature and salinity, but are generally higher at lower temperatures (Fig. A3.1b). For typical

TS values in the central Greenland Sea, the mean root mean square difference is 0.0044 $kg/m^3$ (Fig. A3.1a). At $\sigma_{0.5} = 30.444$ $kg/m^3$, which is taken to represent the limit between intermediate and dense water masses in EOS-80 (Rudels et al., 2002), the density is on average 0.0047 $kg/m^3$ higher when applying the TEOS-10 standard. We therefore consider a density of $\sigma_{0.5} = 30.448$ $kg/m^3$ as the lower limit for GSDW following the TEOS-10 standard.

A4. Reference profile for bottom-reaching convection

In February 1971, a particularly weakly stratified hydrographic profile was obtained in the central Greenland Sea (Malmberg, 1983). This profile, later referred to as the “Malmberg profile”, serves as an example for the occurrence of bottom-reaching convection in the Greenland Sea (e.g., Meincke et al., 1992) and was therefore used as
<figure><img src="image_23.png" /><figcaption>Fig. A2.5. Standard deviation of Conservative Temperature derived from gridding with 1000 bootstrap samples. The color scale corresponds to that of density (Fig. A2.3).</figcaption></figure>

<caption>Table A.2</caption>

Accuracies of temperature, salinity, and pressure measurements. This list is not intended to be complete, but serves to convey an impression of the development of measurement accuracy over time.

<table><thead><tr><th>Accuracy</th><th>Instrument or method</th><th>Period of use</th><th>Reference</th></tr></thead><tbody><tr><td>±3 m</td><td>piezometre</td><td>since the 1870s</td><td>Mohn (1887)</td></tr><tr><td>±0.35 % of full scale range</td><td>strain gauge bridge on CTD</td><td>since the 1980s</td><td>Driemel et al. (2017)</td></tr><tr><td>±6.5 dbar</td><td>strain gauge bridge on CTD</td><td>since the 1980s</td><td>Driemel et al. (2017)</td></tr><tr><td>±0.01 % of full scale range</td><td>precision-machined Si on CTD</td><td>since the 1990s</td><td>Driemel et al. (2017)</td></tr><tr><td>±0.015 % of full scale range</td><td>quartz crystal resonators on CTD</td><td>recent decades</td><td>Driemel et al. (2017)</td></tr><tr><td>±0.3 dbar</td><td>quartz crystal resonators on CTD</td><td>recent decade</td><td>Renfrew et al. (2023)</td></tr><tr><td>±0.01 ‰, with offsets of up to ±0.03 ‰</td><td>titration of water samples</td><td>until the 1970s</td><td>Helland-Hansen and Nansen (1909)</td></tr><tr><td>±0.002 psu</td><td>conductivity of sea water</td><td>since the late 1970s</td><td>(Dickson et al., 1996)</td></tr><tr><td>±0.05 psu</td><td>conductivity sensor on xCTD</td><td>since the late 1970s</td><td>(Mizuno and Watanabe, 1998)</td></tr><tr><td>±0.001 psu - ±0.003 psu</td><td>conductivity sensor on CTD</td><td>recent decades</td><td>(Ullman and Hebert, 2014; Krahmann et al., 2009)</td></tr><tr><td>±0.05 psu - ±0.009 psu</td><td>conductivity sensor on xCTD</td><td>recent decades</td><td>(Renfrew et al., 2023)</td></tr><tr><td>±0.01 °C, up to ±0.05 °C</td><td>reversing thermometer</td><td>until the early 1980s</td><td>Helland-Hansen and Nansen (1909), Aagaard (1968), Dickson et al. (1996)</td></tr><tr><td>±0.001 °C - ±0.002 °C</td><td>temperature sensor on CTD</td><td>recent decades</td><td>Ullman and Hebert (2014), Krahmann et al. (2009)</td></tr><tr><td>±0.05 °C</td><td>temperature sensor on xCTD</td><td>1990s</td><td>Mizuno and Watanabe (1998)</td></tr><tr><td>±0.001 °C</td><td>temperature sensor on xCTD</td><td>recent decade</td><td>Renfrew et al. (2023)</td></tr></tbody></table>
<figure><img src="image_24.png" /><figcaption>**Fig. A3.1.** Density difference between the EOS-80 and TEOS-10 standards. (a) Difference between EOS-80 and TEOS-10 densities for the entire data set. (b) Difference between EOS-80 and TEOS-10 densities for the *T* and *S* range covered by the data set. The contour lines indicate σ<sub>0.5,EOS-80</sub>.</figcaption></figure>

a reference to identify bottom-reaching convection in our data set. Two other profiles obtained from the central Greenland Sea within the same 24 h period do not show similar evidence of bottom-reaching convection (Fig. A4.1).

All profiles indicating convection deeper than 3000 m are shown in Fig. A4.2. All of the profiles have negligible stratification below 1000 m, which is a strong indication of recent convection. As restratification occurs more quickly near the surface, the hydrographic properties of the upper water column may differ from those at depth. Different degrees of restratification can be distinguished among the weakly-stratified profiles, this may be considered a measure of the time that

has passed since the last deep-convection event. Profiles with nearly constant properties throughout the entire water column (red) and profiles with a very shallow surface layer (blue) in Fig. A4.2 suggest that bottom-reaching convection occurred merely days before the profile was recorded (Boccaletti et al., 2007). Some profiles (gray) show pronounced restratification within the upper 1000 m and the week stratification below may thus be a remnant of convection in the previous season. As the amount of late-winter data prior to 1985 is sparse, these summertime profiles showing indirect evidence of bottom-reaching convection are also important for determining when that phase ended (Fig. 4d).
<figure><img src="image_25.png" /><figcaption>**Fig. A4.1.** Profiles from the February 1971 cruise in the central Greenland Sea. The locations of the profiles are marked in the inset (d). The dark-blue profile is the so-called "Malmberg profile".</figcaption></figure>

<figure><img src="image_26.png" /><figcaption>**Fig. A4.2.** Profiles indicating convection deeper than 3000m identified by manual detection. The different colors and line widths were used to visualize the differences in degree of restratification. The thick, black profile is the Malmberg profile used as reference, the red profiles are weakly stratified throughout the water column, the blue profiles have a shallow, restratified surface layer, while the gray profiles show a higher degree of restratification in the upper part of the water column. The locations of the profiles are marked in the inset (d).</figcaption></figure>

A5. Width of the marginal ice zone

The width of the marginal ice zone varied along with the sea-ice patterns. Both the *Odden* and *large ice extent* patterns have large standard deviations in thickness with particularly wide marginal ice zones in the central Greenland Sea (Fig. A5.1). For the *open-ocean* pattern, the marginal ice zone was considerably narrower and more evenly distributed (Fig. A5.1).

A6. Further details on heat-flux distributions

To account for differences in the temporal distribution of heat loss through the winter, we included the full six-month period from November to April in the calculations of the total winter heat loss. In some years, particularly during the *large ice extent* and *Odden* periods, more than 50% of the total heat loss took place in November and December (Fig. A6.1). Toward the end of the 70-year period, under *open-ocean*
<figure><img src="image_27.png" /><figcaption>**Fig. A5.1.** Distribution of width of the marginal ice zone for the three different sea-ice patterns (Fig. 11). The median and standard deviation for each sea-ice pattern are printed in color.</figcaption></figure>
<figure><img src="image_28.png" /><figcaption>Fig. A6.1. Stacked contributions from each winter month to the total winter heat loss. Only the ice-free portion of the Greenland Sea (sea-ice concentration < 15%) was included in the calculation.</figcaption></figure>

<figure><img src="image_29.png" /><figcaption>Fig. A6.2. Characteristics of extreme heat-flux events. The dark-blue dots show the 95th percentile of heat fluxes for each winter (November to April), the dark-blue linear regression line indicates the trend over the 70-year period, and the light-blue line marks the 95th percentile of winter heat fluxes throughout the entire 70-year period. The bars indicate the total heat loss during the extreme events integrated over the ice-free area impacted by each event, color-coded by the total ice-free area. The red line is the linear regression for the total heat loss during extreme events.</figcaption></figure>

conditions, the heat loss was more evenly distributed through the winter.

A large portion of the total winter heat loss takes place during extreme events (Papritz and Spengler, 2017). Over the 70-year period, the intensity of the extreme events declined (Fig. A6.2). However, as the open-ocean area impacted by the extreme events increased, the total heat lost during extreme events increased with decreasing sea-ice extent.

## Appendix B. Supplementary material

Supplementary material related to this article can be found online at https://doi.org/10.1016/j.pocean.2024.103304.

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