After more than a decade of shallow convection, deep convection returned to
the Irminger Sea in 2008 and occurred several times since then to reach
exceptional convection depths (> 1500 m) in 2015 and 2016.
Additionally, deep mixed layers deeper than 1600 m were also reported
southeast of Cape Farewell in 2015. In this context, we used Argo data to
show that deep convection occurred southeast of Cape Farewell (SECF) in 2016
and persisted during two additional years in 2017 and 2018 with a maximum
convection depth deeper than 1300 m. In this article, we investigate the
respective roles of air–sea buoyancy flux and preconditioning of the water
column (ocean interior buoyancy content) to explain this 4-year persistence
of deep convection SECF. We analyzed the respective contributions of the
heat and freshwater components. Contrary to the very negative air–sea
buoyancy flux that was observed during winter 2015, the buoyancy fluxes over
the SECF region during the winters of 2016, 2017 and 2018 were close to the
climatological average. We estimated the preconditioning of the water column
as the buoyancy that needs to be removed (B) from the end-of-summer water
column to homogenize it down to a given depth. B was lower for the winters of 2016–2018 than for the 2008–2015 winter mean, especially due to a vanishing
stratification from 600 down to ∼1300 m. This means that
less air–sea buoyancy loss was necessary to reach a given convection depth
than in the mean, and once convection reached 600 m little additional
buoyancy loss was needed to homogenize the water column down to 1300 m. We
show that the decrease in B was due to the combined effects of the local
cooling of the intermediate water (200–800 m) and the advection of a
negative S anomaly in the 1200–1400 m layer. This favorable
preconditioning permitted the very deep convection observed in 2016–2018
despite the atmospheric forcing being close to the climatological average.
Introduction
Deep convection is the result of a process by which surface waters lose
buoyancy due to atmospheric forcing and sink into the interior of the ocean.
It occurs only where specific conditions are met, including large air–sea
buoyancy loss and favorable preconditioning (i.e., low stratification of the
water column) (Marshall and Schott, 1999). In the
subpolar North Atlantic (SPNA), deep convection takes place in the Labrador
Sea, south of Cape Farewell and in the Irminger Sea
(Kieke and
Yashayaev, 2015; Pickart et al.,
2003; Piron et al., 2017). Deep convection connects the upper and lower
limbs of the Meridional Overturning Circulation (MOC) and transfers climate
change signals from the surface to the ocean interior.
Observing deep convection is difficult because it happens on short timescales and
small spatial scales and during periods of severe weather conditions
(Marshall and Schott, 1999). The onset of the Argo
program at the beginning of the 2000s has considerably increased the number
of available oceanographic data throughout the year. Although the sampling
characteristics of Argo are not adequate to observe the small scales
associated with the convection process itself, Argo data allow for the
description of the overall intensity of the event and the characterization
of the properties of the water masses formed in the winter mixed layer as
well (e.g., Yashayaev and Loder, 2017).
In the Labrador Sea, deep convection occurs every year, yet with different
intensity (e.g., Yashayaev and Clarke, 2008;
Kieke and Yashayaev, 2015). In the Irminger
Sea, Argo and mooring data showed that convection deeper than 700 m happened
during the winters of 2008, 2009, 2012, 2015 and 2016 (Väge et al., 2009;
de Jong et al., 2012, 2018; Piron et al., 2016, 2017;
de Jong and de Steur, 2016; Fröb et al., 2016). Moreover, in winter
2015, deep convection was also observed south of Cape Farewell (Piron et
al., 2017). Excluding winter 2009 when the deep convection event was made
possible thanks to a favorable preconditioning (de Jong et al., 2012), all
events coincided with strong atmospheric forcing (air–sea heat loss). Prior
to 2008, only few deep convection events were reported because the
mechanisms leading to them were not favorable (Centurioni and Gould,
2004) or because the observing system was not adequate
(Bacon, 1997;
Pickart et al., 2003). Nevertheless, the hydrographic properties from the
1990s suggest that deep convection reached as deep as 1500 m in the
Irminger Sea during the winters of 1994 and 1995 (Pickart et al., 2003) and as
deep as 1000 m south of Cape Farewell during winter 1997 (Bacon et al.,
2003).
Positions of all Argo floats north of 55∘ N in
the Atlantic between 1 January and 30 April (a) 2015, (b) 2016, (c) 2017 and (d) 2018 (black and colored points). The colored points and color bar indicate
the mixed layer depth (MLD) when the MLD was deeper than 700 m. The pink circles
indicate the position of the maximal MLD observed SECF each winter. The pink
and cyan boxes delimit the regions used for estimating the time series of
atmospheric forcing and the vertical profiles of buoyancy to be removed in
the SECF region and the Labrador Sea, respectively (SECF: 56.5–59.3∘ N and 45.0–38.0∘ W, Labrador
Sea: 56.5–59.2∘ N and 56–48∘ W).
The convection depths that were reached in the Irminger Sea and south of
Cape Farewell at the end of winter 2015 were the deepest observed in these
regions since the beginning of the 21st century (de Jong et al., 2016;
Piron et al., 2017, Fröb et al., 2016). In this work, we show that deep
convection also happened in a region between south of Cape Farewell and the
Irminger Sea (the pink box in Fig. 1) every winter from 2016 to 2018.
Hereinafter, we will refer to this region as southeast Cape Farewell (SECF).
We investigated the respective role of atmospheric forcing (air–sea buoyancy
flux) and preconditioning (ocean interior buoyancy content) in setting the
convection intensity. We also disentangled the relative contribution of
salinity and temperature anomalies to the preconditioning. The paper is
organized as follows. The data are described in Sect. 2. The methodology is
explained in Sect. 3. We present our results in Sect. 4 and discuss them in
Sect. 5. Conclusions are listed in Sect. 6.
Data
We used temperature (T), salinity (S) and pressure (P) data measured by Argo
floats north of 55∘ N in the Atlantic Ocean. These data were
collected by the International Argo Program (http://www.argo.ucsd.edu/, last access: 9 January 2020;
http://www.jcommops.org/, last access: 9 January 2020) and downloaded from the Coriolis Data
Center (http://www.coriolis.eu.org/, last access: 9 January 2020). Only data flagged as good
(quality control < 3; Argo Data Management Team, 2017) were
considered in our analysis. Potential temperature (θ), density
(ρ) and potential density anomaly referenced to the surface and 1000 dbar (σ0 and σ1, respectively) were estimated from
T, S and P data using TEOS-10 (http://www.teos-10.org/, last access: 9 January 2020).
We used two different gridded products of ocean T and S: ISAS and EN4. ISAS
(Gaillard et al., 2016; Kolodziejczyk et al., 2017) is produced by optimal
interpolation of in situ data. It provides monthly fields at 152 depth levels and at
0.5∘ resolution from 2002 to 2015. Near-real-time data are also
available for 2016 and 2018. EN4 (Good et al., 2013)
is an optimal interpolation of in situ data; it provides monthly T and S at
1∘ spatial resolution and at 42 depth levels for the period 1900
to present.
Net air–sea heat flux (Q, the sum of radiative and turbulent fluxes),
evaporation (E), precipitation (P), wind stress (τx and τy) and sea surface temperature (SST) data were obtained from
the ERA-Interim reanalysis
(Dee
et al., 2011). ERA-Interim provides data with a time resolution of 12 h and a
spatial resolution of 0.75∘. The air–sea freshwater
flux (FWF) was estimated as E–P.
We used monthly absolute dynamic topography (ADT), which was computed from
the daily 0.25∘ resolution ADT data provided by CMEMS (Copernicus
Marine and Environment Monitoring Service, http://www.marine.copernicus.eu, last access: 9 January 2020).
MethodsQuantification of deep convection
We characterized the convection in the SPNA in the winters of 2015–2018 by
estimating the mixed layer depths (MLDs) for all Argo profiles collected in
the SPNA north of 55∘ N from 1 January to 30 April
of each year (Fig. 1). The MLD was estimated as the shallowest of the three
MLD estimates obtained by applying the threshold method
of de Boyer Montégut et al. (2004) to
θ, S and ρ profiles separately. The threshold method computes
the MLD as the depth at which the difference between the surface (30 m) and
deeper levels in a given property is equal to a given threshold. In the case
that visual inspection of the winter profiles showed a thin stratified layer at
the surface, a slightly deeper level (< 150 m) was considered the surface reference level. Following Piron et al. (2017), this threshold was
taken as equal to 0.01 kg m-3 for ρ. For θ and S, we
selected thresholds of 0.1∘ C and 0.012, respectively, because they
correspond to the threshold of 0.01 kg m-3 in ρ. The latter was
previously shown to perform well in the subpolar gyre on density profiles
(Piron et al., 2016). The criteria for temperature and salinity were chosen
to perform well when temperature and salinity anomalies within the
density-defined mixed layer are density-compensated. Our MLD estimates are
comparable to those obtained using MLD determination based on the Pickart et al. (2002) methods (see Sect. S1 and Figs. S1 and S2 in the Supplement).
In this paper, deep convection is characterized by profiles with an MLD deeper
than 700 m (colored points in Fig. 1) because it is the minimum depth that
should be reached for renewing Labrador Sea Water (LSW)
(Yashayaev et al., 2007;
Piron et al., 2016). The winter MLD and
the associated θ, S and ρ properties were examined for the
Labrador Sea and the SECF region by considering the profiles inside the cyan
and pink boxes in Fig. 1, respectively. Those two boxes were defined to
include all Argo profiles with an MLD deeper than 700 m during 2016–2018 and
the minimum of the monthly ADT for either the SECF region or the Labrador
Sea. No deep MLD was recorded in the northernmost part of the Irminger Sea
during this period. We computed the maximum MLD and the MLD third quartile
(Q3) from profiles with an MLD greater than 700 m in each of the two boxes
separately. Q3 is the MLD value that is exceeded by 25 % of the
profiles and is equivalent to the aggregate maximum depth of convection
defined by Yashayaev and Loder (2016). Hereafter, we refer to
Q3 as the aggregate maximum depth of convection. The properties (ρ, θ and S) of the mixed layers were defined for each winter as the
vertical mean from 200 m to the MLD of all profiles with an MLD deeper than 700 m. For further use, we define the deep convection period as follows. For a
given winter, the deep convection period begins the day when the first
profile with a deep (> 700 m) mixed layer is detected and ends the
day of the last detection of a deep mixed layer.
Time series of atmospheric forcing
The air–sea buoyancy flux (Bsurf) was calculated as the sum of the
contributions of Q and FWF (Gill, 1982; Billheimer and
Talley, 2013):
Bsurf=αgρ0cpQ-βgSSS FWF,
where α and β are the coefficients of thermal and saline
expansions, respectively, estimated from surface T and S. The gravitational
acceleration g is equal to 9.8 m s-2, the reference density of sea
water ρ0 is equal to 1026 kg m-3 and the heat capacity of sea
water Cp is equal to 3990 J kg -1∘C-1. SSS is
the sea surface salinity (Q: W m-2, FWF: m s-1).
For easy comparison with previous results, which only considered the heat
component of the buoyancy air–sea flux (e.g.,
Yashayaev and Loder, 2017;
Piron et al., 2017;
Rhein et al., 2017), Bsurf (m2 s-3) was
converted (W m-2) following Eq. (2) and noted Bsurf∗:
Bsurf∗=ρ0cpgαBsurf.
The FWF was also converted (W m-2) using
FWF*=FWFβSSSρ0cpα.
We also computed the horizontal Ekman buoyancy flux (BFek), which can
be decomposed into the horizontal Ekman heat flux (HFek) and salt flux
(SFek).
4BFek=-g(Ue∂xSSD+Ve∂ySSD)Cpα.g5HFek=-(Ue∂xSST+Ve∂ySST)ρ0Cp6SFek=-(Ue∂xSSS+Ve∂ySSS)βρ0Cpα
BFek=SFek – HFek. Ue and Ve are the eastward
and northward components of the Ekman horizontal transport estimated from
the wind stress meridional and zonal components. SSD, SST and SSS are ρ, T and S at the surface of the ocean (BFek, HFek and SFe: J s-1 m-2). Because ERA-Interim does not supply SSD or SSS,
they were estimated from EN4 as follows. The monthly T and S data at 5 m
of depth from EN4 were interpolated on the same time and space grid as the
air–sea fluxes from ERA-Interim (12 h and 0.75∘, respectively).
SSD was estimated from those interpolated EN4 data (SST and SSS). Properties
at 5 m of depth were considered to be representative of the Ekman layer. Data
at locations where the ocean bottom was shallower than 1000 m were excluded from
the analysis to avoid regions covered by sea ice.
Following Piron et al. (2016), the
time series of atmospheric forcing were estimated for the SECF region and
the Labrador Sea as follows. First, the gridded air–sea flux data and the
horizontal Ekman fluxes were averaged over the pink (SECF region) and cyan
(Labrador Sea) boxes (Fig. 1). Second, we estimated the accumulated fluxes
from 1 September to 31 August the year after. Finally, we computed
the time series of the anomalies of the accumulated fluxes from 1 September
to 31 August with respect to the 1993–2016 mean.
Finally, in order to quantify the net intensity of the atmospheric forcing
over the winter, we computed estimates of Bsurf*+BFek fluxes
accumulated from 1 September to 31 March the year after. Following Piron et
al. (2017), the associated errors were calculated by a Monte Carlo
simulation using 50 random perturbations of Q, FWF and Bsurf. The error
amounted to 0.05, 0.04 and 0.03 J m-2 for Bsurf*, Q and FWF*,
respectively. The error of the horizontal Ekman buoyancy transport was also
estimated by a Monte Carlo simulation and amounted to 0.04 J m-2.
Preconditioning of the water column
The preconditioning of the water column was evaluated as the buoyancy that
has to be removed (B(zi)) from the late summer density profile to
homogenize it down to a depth zi:
Bzi=gρ0σ0zizi-gρ0∫zioσ0(z)dz,
where σ0(z) is the vertical profile of potential density anomaly
estimated from the profiles of T and S measured by Argo floats in September
in the given region (pink or cyan box in Fig. 1).
Following Schmidt and Send (2007), we split B into a
temperature (Bθ) and salinity (BS) term:
8Bθzi=-(gαθzizi-gα∫zioθ(z)dz),9BSzi=gβSzizi-gβ∫zioS(z)dz.
In order to compare the preconditioning with the heat to be removed and/or
air–sea heat fluxes, B, Bθ and BS are reported in Joules per square meter. B, Bθ and BS were estimated for a given
year from the mean of all September profiles of B, Bθ and
BS. The associated errors were estimated as std(B)/√n, where
n is the number of profiles used to compute the September mean values.
Properties of the deep convection SECF and in the Labrador Sea in
winters 2015–2018. We show: the maximal MLD observed, the aggregate
maximum depth of convection, the σ0, S and θ of the
winter mixed layer formed during the convection event and n, which is the
number of Argo profiles indicating deep convection. The uncertainties given
with σ0, S and θ are the standard deviation of the n
values considered to estimate the mean values.
SECFMaximal MLD (m)Aggregate max. depthσ0Salinityθnof convection (m)W20151710120527.733±0.00734.866±0.0133.478±0.13029W20161575147127.746±0.00234.871±0.0033.388±0.0323W20171400125127.745±0.00734.868±0.0073.364±0.1093W20181300130027.748±0.00134.859±0.0033.263±0.0312Labrador SeaMaximal MLDAggregate max. depthσ0Salinityθnof convection (m)W20151675150427.733±0.00934.842±0.0103.279±0.03641W20161801162027.743±0.00634.836±0.0103.124±0.04718W20171780167427.752±0.00834.853±0.0093.172±0.02926W20182020186627.756±0.00634.855±0.0103.145±0.08313ResultsIntensity of deep convection and properties of newly formed LSW
We examine the time evolution of the winter mixed layer SECF since the
exceptional convection event of winter 2015 (W2015 hereinafter) (Table 1 and
Figs. 1–3). In W2015, we recorded a maximum MLD of 1710 m south of Cape
Farewell (Fig. 1a), in line with Piron et al. (2017). The maximum MLD of
1575 m observed for W2016 (Fig. 1b) is compatible with the active mixed
layer > 1500 m observed in a mooring array in the central
Irminger Sea by de Jong et al. (2018). For W2015 and W2016, the aggregate
maximum depth of convection was 1205 and 1471 m, respectively (Table 1).
In W2017, deep convection was observed from three Argo profiles (Figs. 1c and 2a–c). The maximum MLD of 1400 m was observed on 16 March 2017
at 56.65∘ N–42.30∘ W. In W2018, the maximum MLD of
1300 m was observed on 24 February at 58.12∘ N, 41.84∘ W (Figs. 1d, 2d–f). Float 5903102 measured an MLD of 1100 m south of Cape
Farewell (Fig. 1d), but the estimated MLDs coincided with the deepest levels
of measurement of the float so that these estimates, possibly biased low
(see Fig. 2d–f), were discarded from our analysis. These results show that
convection deeper than 1300 m occurred during four consecutive winters
SECF.
Vertical distribution of σ0, S and θ
of Argo profiles showing an MLD deeper than 700 m SECF in winter 2017 (a, b, c) and in winter 2018 (d, e, f). The black points indicate the MLD. The
triangles in (d) are the MLD that coincided with the maximal profiling
pressure reached by the float. In the legend, the float and cycle of each
profile are indicated.
Although the number of floats showing deep convection in W2017 and W2018 was
small (three and two floats, respectively), it represented a significant percentage of the
floats operating in the SECF box at that time. The percentage of floats
showing deep convection in the SECF region was computed for the deep
convection periods defined from 15 January to 21 April 2015, 22 February to 21 March 2016, 16 March to 4 April 2017 and 24 February to 26 March 2018. The longest period of deep convection
occurred in W2015 and the shortest in 2017. The percentages of floats showing
deep convection during the deep convection period are 73 %, 50 %, 33 % and 50 % for the winters of 2015, 2016, 2017 and 2018, respectively. The
lowest percentage is found for W2017, but it is still substantial. It might
reflect the fact that for this specific year floats showing a deep MLD were found in
the southwestern corner of the SECF box only, suggesting that convection did
not occur over the full box.
TS diagrams in the mixed layer for profiles with an MLD
deeper than 700 m during the winters of 2015, 2016, 2017 and 2018 for (a) the
Labrador Sea and (b) SECF. The properties of the mixed layers were estimated
as the vertical means between 200 m and the MLD.
The properties (σ0, S and θ) of the end-of-winter mixed
layer were estimated for the four winters (Table 1 and Fig. 3). We observed
that, between W2015 and W2018, the water mass formed by deep convection
significantly densified and cooled by 0.019 kg m-3 and 0.215 ∘C, respectively (see Table 1 and Fig. 3).
In the Labrador Sea, the aggregate maximum depth of convection increased
from 2015 to 2018 (see Table 1). Deep convection observed in the Labrador
Sea in W2018 was the most intense since the beginning of the Argo era (see
Fig. 2c in Yashayaev and Loder, 2016). From W2015 to W2018,
newly formed LSW cooled, became saltier, and densified by 0.134 ∘C,
0.013 and 0.023 kg m-3, respectively (Table 1).
The water mass formed SECF is warmer and saltier than that formed in the
Labrador Sea (Fig. 3). The deep convection SECF is always shallower than in
the Labrador Sea. This result is discussed later in Sect. 5.
Analysis of the atmospheric forcing southeast of Cape Farewell
The seasonal cycles of Bsurf* and Q are in phase and of the same order
of magnitude, while FWF*, which is positive and 1 order of magnitude lower
than Q, does not present a seasonal cycle (Fig. S3). The means (1993–2018) of the cumulative sums from 1 September to 31 March of Q, FWF* and
Bsurf* estimated over the SECF box (Fig. 1) are -2.46±0.43×109, 0.28±0.10×109 and -2.22±0.49×109 J m-2, respectively. Bsurf* is 10 %
lower on average than Q because of the buoyancy addition by FWF*.
Considering the Ekman transports, the 1993–2018 means of the accumulated
BFek, HFek and SFek from 1 September to 31 March amount to
0.37±1.15×108, -0.35±1.36×108 and 0.02±2.04×108×109 J m-2,
respectively. The horizontal Ekman heat flux is negative, while the Ekman
buoyancy flux is positive. This buoyancy gain indicates a southeastward
transport of surface freshwater caused by dominant winds from the southwest.
It is noteworthy that BFek is 1 order of magnitude smaller than the
Bsurf*.
Time series of anomalies of accumulated (a)Bsurf*, (b)Q,
(c) FWF*(d) BFek, (e) HFek and (f) SFek averaged in the SECF
region. They are anomalies with respect to 1993–2016. The accumulation
was from 1 September to 31 August the following year. The winter North Atlantic Oscillation (NAO) index
(Hurrell et al., 2018) is also represented in (g). The gray line in (a) is the sum
of the anomalies of accumulated Bsurf* and BFek. Note that the
range of values in the y axis is not the same in all the plots.
The total atmospheric forcing SECF was quantified as the sum of Bsurf*
and BFek. The anomalies of accumulated fluxes from 1 September to 31 August the year after, with respect to the mean 1993–2016, are displayed
in Fig. 4 for the SECF box. The gray line in Fig. 4a is the total
atmospheric forcing anomaly (Bsurf* plus BFek). We identify years
with very negative buoyancy loss in the SECF region, e.g., 1994, 1999, 2008,
2012 and 2015. The very negative anomalies of atmospheric forcing in 1999
and 2015 were caused by the very negative anomalies in both Bsurf*
(Fig. 4a) and BFek (Fig. 4d). This correlation was not observed for all
the years presenting a negative anomaly of atmospheric forcing. It is noteworthy that
during W2016, W2017 and W2018, the anomaly of atmospheric forcing was close
to zero.
Contrary to the very negative anomaly in atmospheric fluxes over the SECF
region observed for W2015, the atmospheric fluxes were close to the mean
during W2016, W2017 and W2018.
Time evolutions of vertical profiles measured from Argo
floats in the SECF region: (a)θ, (b)S, (c)σ1 and (d) the thickness of 0.01 kg m-3 thick σ1 layers. The white
horizontal bars in plots (a), (b) and (c) indicate the maximal convection depth
observed in the Irminger Sea or SECF when deep convection occurred. The white
line in plot (a) indicates the depth of the isotherm 3.6 ∘C. The
black vertical ticks on the x axes of plot (b) indicate times of Argo
measurements. These figures were created from all Argo profiles reaching
deeper than 1000 m in the SECF region (56.5–59.3∘ N,
45–38∘ W; pink box in Fig. 1). The yearly numbers of
Argo profiles used in this figure are shown in Fig. S5.
Analysis of the preconditioning of the water column southeast of Cape Farewell
Our hypothesis is that the exceptional deep convection that happened in
W2015 in the SECF region favorably preconditioned the water column for deep
convection the following winters. The time evolutions of θ, S,
σ1 and Δσ1=0.01 kg m-3 layer
thicknesses (Fig. 5) show a marked change in the hydrographic properties of
the SECF region at the beginning of 2015 caused by the exceptional deep
convection that occurred during W2015 (see Piron et al., 2017). The
intermediate waters (500–1000 m) became colder than the years before,
and despite a slight decrease in salinity, the cooling caused the density
to increase (Fig. 5c). Figure 5d shows Δσ1=0.01 kg m-3 layer thicknesses larger than 600 m appearing at the end of W2015
for the first time since 2002. In the density range 32.36–32.39 kg m-3, these layers remained thicker than ∼450 m during
W2016 to W2018. This indicates low stratification at intermediate depths and
a favorable preconditioning of intermediate waters for deep convection
initiated by W2015 deep convection. The denser density of the core of the
thick layers in 2017–2018 compared with 2015–2016 agrees with the
densification of the mixed layer SECF shown in Table 1 and Fig. 3.
Vertical profile of (a) the buoyancy to be removed (B), (b) the thermal component (Bθ) and (c) the salinity component
(BS). They were calculated from all Argo data measured in the SECF box
(see Fig. 1) in September before the winter indicated in the legend. For
W2015 and W2018, we considered data from 15 August to 30 September 2017 because
not enough data were available in September. The number of Argo profiles
taken into account to estimate the B profiles was more than 10 for all the
winters.
B(zi) is our estimate of the preconditioning of the water column before
winter (see the Methods section). Figure 6a shows that, deeper than 100 m, B for W2016,
W2017 and W2018 was smaller than B for W2015 or B for the mean W2008–W2014. Furthermore, for W2016, W2017 and 2018, B remained nearly constant
with depth between 600 and 1300 m, which means that once the water column
has been homogenized down to 600 m, little additional buoyancy loss results
in the homogenization of the water column down to 1300 m. Both conditions,
(i) less buoyancy to be removed and (ii) the absence of a gradient in the B
profile down to 1300 m, indicate a more favorable preconditioning of the
water column for W2016, W2017 and W2018 than during W2008–W2015.
To understand the relative contributions of θ and S to the
preconditioning, we computed the thermal (Bθ) and haline
(BS) components of B (B=Bθ+BS). In general,
Bθ (BS) increases with depth when θ decreases (S
increases) with depth. On the contrary, a negative slope in a Bθ
(BS) profile corresponds to θ increasing (S decreasing) with
depth and is indicative of a destabilizing effect. The negative slopes in
Bθ and BS profiles are not observed simultaneously because
density profiles are stable.
We describe the relative contributions of Bθ and BS to B by
looking first at the mean 2008–2014 profiles (discontinuous blue lines in
Fig. 6). Bθ accounts for most of the increase in B from the
surface to 800 m and below 1400 m (see Fig. 6a and b). The negative
slope in the BS profile between 800 and 1000 m (Fig. 6c) slightly
reduces B (Fig. 6a) and is due to the decrease in S associated with the core
of LSW (see Fig. 3 in Piron et al., 2016). In the layer 1000–1400 m,
the increase in B (Fig. 6a) is mainly explained by the increase in BS (Fig. 6c), which follows the increase in S in the transition from LSW to
Iceland Scotland Overflow Water (ISOW). This transition layer will be
referred to hereinafter as the deep halocline. The
preconditioning of the water column is usually analyzed in terms of heat
(e.g., Piron et al., 2015, 2017). The decomposition of B in Bθ
and BS reveals that θ governs B in the layer 0–800 m. S
tends to reduce the stabilizing effect of θ in the layer 800–1000 m and reinforces it in the layer 1000–1400 m by adding up to 1×109 J m2 to B.
Decomposition of the profiles of buoyancy to be removed (B,
continuous lines) in its thermal (Bθ, dotted lines) and salinity
(BS, dashed lines) components in (a) the SECF region; (b) the Labrador
Sea. The BS components for W2016 and W2018 were added to show the
evolution of the depth of the deep halocline.
In order to further understand why the SECF region was favorably
preconditioned during the winters of 2016–2018, we compare the Bθ and
BS of W2017, which was the most favorably preconditioned winter, with
the mean 2008–2014 (Fig. 7a). From the surface to 1600 m, Bθ
and BS were smaller for W2017 than for the mean 2008–2014. There are
two additional remarkable features. First, in the layer 500–1000 m, the
large reduction of Bθ compared to the 2008–2014 mean mostly
explains the decrease in B in this layer. Second, the more negative value of
BS in the layer 1100–1300 m, compared to the 2008–2014 mean,
eroded the Bθ slope, making the B profile more vertical for W2017
than for the mean. The more negative contribution of BS in the layer
1100–1300 m comes from the fact that the deep halocline was deeper for
W2017 (1300 m; see orange dashed line in Fig. 7a) than for the mean 2008–2014 (1000 m; see blue dashed line in Fig. 7a). Finally, we note that the
profiles of B(zi), Bθ(zi) and BS(zi) for
W2016 and W2018 are more similar to the profiles of W2017 than to those of
W2015 or to the mean 2008–2014 (see Fig. 6), which indicates that the
water column was also favorably preconditioned for deep convection in W2016
and W2018 for the same reasons as in W2017.
Evolution of vertical profiles of monthly anomalies of (a)σ0, (b)θ and (c)S at 58∘ N, 40∘ W.
The anomalies were estimated from the ISAS database
(Gaillard et al., 2016) and were referenced to the monthly
mean estimated for 2002–2016.
The origin of the changes in B is now discussed from the time evolutions of
the monthly anomalies of θ, S and σ0 at 58∘ N–40∘ W, which is at the center of the SECF box (Fig. 8). The time
evolutions there are similar to those at any other location inside the SECF
box. These anomalies were computed using ISAS Gaillard et
al., 2016) and were referenced to the monthly mean of 2002–2016. A
positive anomaly of σ0 appeared in 2014 between the surface and
600 m (Fig. 8a) and reached 1200 m in 2015 and beyond. This positive anomaly
of σ0 correlates with a negative anomaly of θ. The
latter, however, reached ∼1400 m of depth in 2016, which is
deeper than the positive anomaly of σ0. The negative anomaly of
S between 1000 and 1500 m that appeared in 2015 and strongly reinforced in
2016 caused the negative anomaly in σ0 between 1200 and 1500 m
(the density anomaly caused by the negative anomaly in θ between
1200 and 1400 m does not balance the density anomaly caused by the negative
anomaly of S).
The θ and S anomalies in the water column during 2016–2018
explain the anomalies of B, Bθ and BS and can be summarized
as follows. On the one hand, the properties of the surface waters (down to
500 m) were colder than previous years and, despite the fact that they were also fresher,
they were denser. The density increase in the surface water reduced the
density difference with the deeper-lying waters. The intermediate layer (500–1000 m) was also favorably preconditioned due to the observed cooling.
Additionally, in the layer 1100–1300 m, the large negative anomaly of
BS with respect to its mean is explained by the decrease in S in this
layer, which caused a decrease in σ0 and consequently reduced
the σ0 difference with the shallower-lying water. The decrease
in S also resulted in a deepening of the deep halocline.
Horizontal distribution of the anomalies of S(a, b, c), θ(d, e, f) and σ0(g, h, i) in
the layer 1200–1400 m in February 2015 (a, d, g), February 2016 (b, e, h) and February 2017 (c, f, i). The monthly anomalies
were estimated from the ISAS database and were referenced to the period 2002–2016.
Atmospheric forcing versus preconditioning of the water column
We now use the estimates of the accumulated atmospheric forcing (Bsurf*+BFek) from 1 September to 31 March the year after (see Fig. S4) to
predict the maximum convection depth for a given winter based on September
profiles of B. The predicted convection depth is determined as the depth at
which B(zi) (Fig. 6a) equals the accumulated atmospheric forcing. The
associated error was estimated by propagating the error in the atmospheric
forcing (0.05×109 J m-2). The accumulated atmospheric forcing
amounted to -3.21×109±0.05-2.21±0.04×109, -2.01±0.05×109 and -2.47±0.05×109 J m-2 for W2015, W2016, W2017 and W2018,
respectively. We found predicted convection depths of 1085±20,
1285±20, 1415±20 and 345±20 m for W2015,
W2016, W2017 and W2018, respectively. We consider the aggregate maximum
depth of convection to be the observed estimate of the MLD (Table 1). The
predicted MLD agrees with the observed MLD within ±200 m. The
differences could be due to errors in the atmospheric forcing (Josey et al.,
2018), lateral advection and/or spatial variation in the convection
intensity within the box not captured by the Argo sampling.
The satisfactory predictability of the convection depth with our 1-D model
indicates that deep convection occurred locally. In spite of the fact that the atmospheric
forcing was close to mean (1993–2016) conditions during W2016, W2017 and
W2018, convection depths > 1300 m were reached in the SECF
region. This was only possible thanks to the favorable preconditioning.
Discussion
Deep convection happens in the Irminger Sea and south of Cape Farewell
during specific winters characterized by a strong atmospheric forcing (high
buoyancy loss), a favorable preconditioning (low stratification) or both at
the same time (Bacon et al., 2003;
Pickart et al., 2003). In the
Irminger Sea, strong atmospheric forcing explained, for instance, the very
deep convection (reaching depths greater than 1500 m) observed in the early
1990s (Pickart et al., 2003) and in W2015 (de Jong et al. 2016; Fröb et
al., 2016; Piron et al., 2017). It also explained the return of deep
convection in W2008 (Väge et al., 2009) and in W2012
(Piron et al., 2016). The favorable
preconditioning caused by the densification of the mixed layer during W2008
favored a new deep convection event in W2009 despite neutral atmospheric
forcing (de Jong et al., 2012). Similarly, the preconditioning observed
after W2015 in the SECF region favored deep convection in W2016 (this work).
The favorable preconditioning persisted for three consecutive winters (2016–2018) in the SECF region, which allowed deep convection although atmospheric
forcing was close to the climatological values. Why did this favorable
preconditioning persist in time?
We previously showed that during 2016–2018 two hydrographic anomalies
affected different ranges of the water column in the SECF box: a cooling
intensified in the layer 200–800 m and a freshening intensified in the 1000–1500 m layer. Those resulted in a decrease in the vertical density
gradient between the intermediate and the deeper layers, creating a favorable
preconditioning of the water column. Note that the cooling affected the
layer from the surface to 1400 m, and the freshening affected the layer from
the near surface to 1600 m, but the cooling and the freshening were intensified
at different depth ranges (Fig. 8).
We see in Fig. 5a a sudden decrease in θ in the intermediate layers
in 2015 compared to the previous years. It indicates that the decrease in
θ of the intermediate layer likely originated locally during W2015
when extraordinary deep convection happened. A slight freshening of the
water column (400–1500 m) appeared in 2015, likely caused by the W2015
convection event; then it decreased before a second S anomaly intensified in
2016 between 1100 and 1400 m (Fig. 8c). It is unlikely that this second
anomaly was exclusively locally formed by deep convection because it
intensified during summer 2016. Our hypothesis is that this second S anomaly
originated in the Labrador Sea and was further transferred to the SECF
region by the cyclonic circulation encompassing the Labrador Sea and
Irminger Sea at these depths
(Daniault et
al., 2016; Ollitrault and Colin de
Verdière, 2014; Lavander et al., 2000; Straneo et al., 2003). This is
corroborated by the 2-D evolution of the anomalies in S in the layer 1200–1400 m (Fig. 9): a negative anomaly in S appeared in the Labrador Sea in
February 2015, which was transferred southward and northeastward in February 2016 and intensified over the whole SPNA in February 2017. By this
mechanism, the advection from the Labrador Sea contributed to create
property anomalies in the water column. However, the buoyancy budget showed
that this was a minor contribution compared to the buoyancy loss due to the
local air–sea flux, even if it was essential to preconditioning the water
column for deep convection.
We now compare the atmospheric forcing and the preconditioning of the water
column in the SECF region with those of the nearby Labrador Sea where deep
convection happens almost every year. The atmospheric forcing over the
Labrador Sea is ∼15 % larger than that over the SECF
region: the means (1993–2018) of the atmospheric forcing, defined as the
time-accumulated Bsurf*+BFek from 1 September to 31 March the
year after, are -2.61±0.55×109 J m-2 in the Labrador Sea
and -2.18±0.54×109 J m-2 in the SECF region. The
difference was larger during the period 2016 – 2018 when the atmospheric
forcing equaled -3.10±0.19×109 J m-2 in the Labrador Sea
and -2.23±0.23×109 J m-2 in the SECF region. In terms of
preconditioning, the 2008–2014 mean B profile (blue continuous lines in
Fig. 7) was lower by ∼0.5×109 J m-2 in the
Labrador Sea than SECF for the surface to the 1000 m layer and by more than 1×109 J m-2 below 1200 m. This indicates that the water column was
more favorably preconditioned in the Labrador Sea than in the SECF region
during 2008–2014. Differently, B for W2017 shows slightly lower values
from the surface to 1300 m in the SECF region than in the Labrador Sea (see
orange lines in Fig. 7). However, B in the Labrador Sea remains constant
down to the depth of the deep halocline between LSW and North Atlantic Deep
Water (NADW) at 1700 m. In the SECF region, the deep halocline remained at
∼1300 m between 2016 and 2018 (see BS lines in Fig. 7a). Differently, in the Labrador Sea, the deep halocline deepened from
1200 m for the mean to 1735, 1775 and 1905 m in W2016, W2017 and
W2018, respectively (see the dashed lines in Fig. 7b). The deep halocline acts
as a physical barrier for deep convection in both the SECF region and the
Labrador Sea, but because the deep halocline is deeper in the Labrador Sea
than in the SECF region, the preconditioning is more favorable to a deeper
convection in the Labrador Sea than in the SECF region. Summarizing, in
the winters of 2016–2018 in the Labrador Sea, both atmospheric forcing and
preconditioning of the water column allowed for the deepest convection depth in
the Labrador Sea since the beginning of the Argo period (comparison of our
results with those of Yashayaev and Loader, 2017). In contrast, in the SECF
region during the same period, the atmospheric forcing was close to
climatological values, and the favorable preconditioning of the water column
allowed for 1300 m of depth convection, which was exceptional for the SECF region.
The Labrador Sea, SECF region and Irminger Sea are three distinct deep
convection sites (e.g., Yashayaev et al., 2007; Bacon et al., 2003; Pickart
et al., 2003; Piron et al., 2017). In this work, we give new insights on the
connections between the different sites, showing how the lateral advection of
fresh LSW formed in the Labrador Sea favored the preconditioning in the SECF
region, fostering deeper convection.
Climate models forecast an increasing input of freshwater in the North Atlantic
due to ice melting under present climate change (Bamber et al., 2018), which
could reduce or even shut down the deep convection in the North Atlantic
(Yang et al., 2016;
Brodeau and Koenigk, 2016). We observed a
fresh anomaly in the surface waters in regions close to the eastern coast of
Greenland in 2016 that extended to the whole Irminger Sea in 2017 (Fig. S6).
However, this surface freshening did not hamper the deep convection in the
SECF region, possibly because the surface water also cooled.
Swingedouw et al. (2013) indicated that the freshwater signal due
to Greenland ice sheet melting is mainly accumulating in the Labrador Sea.
However, no negative anomaly of S was detected in the surface waters of the
Labrador Sea (Fig. S6). It might be explained by the intense deep convection
affecting the Labrador Sea since 2014 that could have transferred the
surface freshwater anomaly to the ocean interior. This suggests that, in the
last years, the interactions between expected climate change anomalies and
the natural dynamics of the system combined to favor very deep convection.
This, however, does not foretell the long-term response to climate change.
Conclusions
During 2015–2018 winter deep convection happened in the SECF region,
reaching deeper than 1300 m. The deep convection of W2015 was observed over
a larger region and during a longer period of time than the deep convection
events of the winters of 2016, 2017 and 2018. Despite these differences, it is the
first time that deep convection, with a maximum convection depth greater than
1300 m, was observed in this region during four consecutive winters.
The atmospheric forcing and preconditioning of the water column were
evaluated in terms of buoyancy. We showed that the atmospheric forcing is
10 % weaker when evaluated in terms of buoyancy than in terms of heat
because of the non-negligible effect of the freshwater flux. The analysis of
the preconditioning of the water column in terms of buoyancy to be removed
(B) and its thermal and salinity terms (Bθ and BS) revealed
that Bθ dominated the B profile from the surface to 800 m, and
BS reduced the B in the 800–1000 m layer because of the low salinity of
LSW. Deeper, BS increased B due to the deep halocline (LSW–ISOW) that
acted as a physical barrier limiting the depth of the convection.
During 2016–2018, the air–sea buoyancy losses were close to the
climatological values, and very deep convection was possible thanks to
the favorable preconditioning of the water column. It was surprising that
these events reached convection depths similar to those observed in W2012
and W2015, when the latter were caused by high air–sea buoyancy loss
intensified by the effect of strong wind stress. It was also surprising that
the water column remained favorably preconditioned during three consecutive
winters without strong atmospheric forcing. In this paper, we studied the
reasons why this happened.
The preconditioning for deep convection during 2016–2018 was particularly
favorable due to the combination of two types of hydrographic anomalies
affecting different depth ranges. First, the surface and intermediate waters
(down to 800 m) were favorably preconditioned because buoyancy (density)
decreased (increased) due to the cooling caused by the atmospheric forcing.
Second, buoyancy (density) increased (decreased) in the layer 1200–1400 m due to the decrease in S caused by the lateral advection of fresher LSW
formed in the Labrador Sea. The S anomaly of this layer resulted in a deeper
deep halocline. Hence, the cooling of the intermediate water was essential
to reach a convection depth of 800–1000 m, and the freshening in the layer
1200–1400 m and the associated deepening of the deep halocline allowed for
very deep convection (> 1300 m) in W2016–W2018.
Data availability
The Argo data were collected and made freely available by the International Argo Program and the national programs that contribute to it (10.17882/42182; Argo group, 2019).
The ISAS data were downloaded from the SEANOE site (10.17882/52367; Kolodziejczyk et al., 2017).
The EN4 data are available at http://hadobs.metoffice.com/en4/download.html (Good et al., 2013).
The ERA-Interim data are available at https://www.ecmwf.int/en/forecasts/datasets/reanalysis-datasets/era-interim (Dee et al., 2011).
The NAO data were downloaded from the UCAR Climate Data Guide website (https://climatedataguide.ucar.edu/climate-data/hurrell-north-atlantic-oscillation-nao-index-station-based, last access: 9 January 2020; Hurrell et al., 2018).
The Ssalto/Duacs altimeter products were produced and distributed by the Copernicus Marine and Environment Monitoring Service (CMEMS) (http://marine.copernicus.eu/services-portfolio/access-to-products/?option=com_csw&view=details&product_id=SEALEVEL_GLO_PHY_L4_REP_OBSERVATIONS_008_047, last access: 9 January 2020).
The supplement related to this article is available online at: https://doi.org/10.5194/os-16-99-2020-supplement.
Author contributions
PZ treated and analyzed the data. PZ and HM
interpreted the results. PZ, HM and VT discussed the results and wrote the
paper.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We thank all the navy staff and PIs contributing to the deployment of Argo floats; without them this work would not be possible. We thank Marieke Femke de Jong and two anonymous reviewers; their comments and suggestions have enriched our work.
Financial support
This paper is a contribution to the EQUIPEX NAOS project funded by the French National Research Agency (ANR) under reference ANR-10-EQPX-40.
Review statement
This paper was edited by Neil Wells and reviewed by Femke de Jong and two anonymous referees.
ReferencesArgo Data Management Team: Argo user's manual V3.2,
10.13155/29825, 2017.Argo group: Argo float data and metadata from Global Data Assembly Centre
(Argo GDAC), SEANOE, 10.17882/42182, 2019.Bacon, S.: Circulation and Fluxes in the North Atlantic between Greenland
and Ireland, J. Phys. Oceanogr., 27, 1420–1435, 10.1175/1520-0485(1997)027<1420:CAFITN>2.0.CO;2, 1997.Bacon, S., Gould, W. J., and Jia, Y.: Open-ocean convection in the Irminger
Sea, Geophys. Res. Lett., 30, 1246, 10.1029/2002GL016271, 2003.Bamber, J. L., Tedstone, A. J., King, M. D., Howat, I. M., Enderlin, E. M.,
van den Broeke, M. R., and Noel, B.: Land Ice Freshwater Budget of the
Arctic and North Atlantic Oceans: 1. Data, Methods, and Results,
J. Geophys. Res.-Oceans, 123,
10.1002/2017JC013605, 2018.Billheimer, S. and Talley, L. D.: Near cessation of Eighteen Degree Water
renewal in the western North Atlantic in the warm winter of 2011–2012,
J. Geophys. Res.-Oceans, 118, 6838–6853, 10.1002/2013JC009024, 2013.Brodeau, L. and Koenigk, T.: Extinction of the northern oceanic deep
convection in an ensemble of climate model simulations of the 20th and 21st
centuries, Clim. Dynam., 46, 2863–2882, 10.1007/s00382-015-2736-5,
2016.
Centurioni and Gould, W. J.: Winter conditions in the Irminger Sea observed
with profiling floats, J. Marine Res., 62, 313–336, 2004.Daniault, N., Mercier, H., Lherminier, P., Sarafanov, A., Falina, A.,
Zunino, P., and Gladyshev, S.: The northern North Atlantic Ocean mean
circulation in the early 21st century, Prog. Oceanogr., 146, 142–158,
10.1016/j.pocean.2016.06.007, 2016.de Boyer Montégut, C., Madec, G., Fischer, A. S., Lazar, A., and
Iudicone, D.: Mixed layer depth over the global ocean: An examination of
profile data and a profile-based climatology,
J. Geophys. Res.-Oceans, 109, 1–20,
10.1029/2004JC002378, 2004.de Jong, M. F. and de Steur, L.: Strong winter cooling over the Irminger
Sea in winter 2014–2015, exceptional deep convection, and the emergence of
anomalously low SST, Geophys. Res. Lett., 43, 7106–7113,
10.1002/2016GL069596, 2016.de Jong, M. F., Van Aken, H. M., Våge, K., and Pickart, R. S.:
Convective mixing in the central Irminger Sea: 2002–2010,
Deep-Sea Res. Pt. I, 63, 36–51,
10.1016/j.dsr.2012.01.003, 2012.de Jong, M. F., Oltmanns, M., Karstensen, J., and de Steur, L.: Deep
Convection in the Irminger Sea Observed with a Dense Mooring Array,
Oceanography, 31, 50–59, 10.5670/oceanog.2018.109, 2018.Dee, D. P., Uppala, S. M., Simmons, A. J., Berrisford, P., Poli, P.,
Kobayashi, S., and Vitart, F.: The ERA-Interim reanalysis: Configuration and
performance of the data assimilation system,
Q. J. Roy. Meteor. Soc., 137, 553–597,
10.1002/qj.828, 2011 (data available at: https://www.ecmwf.int/en/forecasts/datasets/reanalysis-datasets/era-interim, last access: 9 January 2020).Fröb, F., Olsen, A., Våge, K., Moore, G. W. K., Yashayaev, I.,
Jeansson, E., and Rajasakaren, B.: Irminger Sea deep convection injects
oxygen and anthropogenic carbon to the ocean interior, Nat. Commun.,
7, 13244, 10.1038/ncomms13244,
2016.Gaillard, F., Reynaud, T., Thierry, V., Kolodziejczyk, N., and Von
Schuckmann, K.: In situ-based reanalysis of the global ocean temperature and
salinity with ISAS: Variability of the heat content and steric height,
J. Climate, 29, 1305–1323, 10.1175/JCLI-D-15-0028.1, 2016.
Gill, A. E.: Atmosphere-Ocean Dynamics, vol. 30, Academic, San Diego,
CA, 1982.Good, S. A., Martin, M. J., and Rayner, N. A.: EN4: Quality controlled
ocean temperature and salinity profiles and monthly objective analyses with
uncertainty estimates, J. Geophys. Res.-Oceans, 118, 6704–6716,
10.1002/2013JC009067, 2013 (data available at: http://hadobs.metoffice.com/en4/download.html, last access: 9 January 2020).Hurrell, J. and National Center for Atmospheric Research Staff (Eds.): The Climate Data Guide: Hurrell North Atlantic Oscillation (NAO) Index (station-based), available at: https://climatedataguide.ucar.edu/climate-data/hurrell-north-atlantic-oscillation-nao-index-station-based (last access: 28 June 2018), 2018.Josey, S. A., Hirschi, J. J.-M., Sinha, B., Duchez, A., Grist, J. P., and Marsh,
R.: The Recent Atlantic Cold Anomaly: Causes, Consequences, and Related
Phenomena, Annu. Rev. Mar. Sci., 10,
475–501, 10.1146/annurev-marine-121916-063102,
2018.Kieke, D. and Yashayaev, I.: Studies of Labrador Sea Water formation and
variability in the subpolar North Atlantic in the light of international
partner and collaboration, Prog. Oceanogr., 132, 220–232,
10.1016/j.pocean.2014.12.010, 2015.Kolodziejczyk, N., Prigent-Mazella, A., and Gaillard, F.: ISAS-15
temperature and salinity gridded fields, Seanoe,
10.17882/52367, 2017.
Lavender, K. L., Davis, R. E., and Owens, W. B.: Mid-depth recirculation
observed in the interior Labrador and Irminger seas by direct velocity
measurements, Nature, 407, 2000.Marshall, J. and Schott, F.: Open-Ocean Convection Theory, and Models
Observations, Rev. Geophys., 37, 1–64, 10.1029/98RG02739, 1999.Ollitrault, M. and Colin de Verdière, A.: The Ocean General
Circulation near 1000-m Depth, J. Phys. Oceanogr., 44, 384–409,
10.1175/JPO-D-13-030.1, 2014.
Pickart, R. S., Torres, D. J., and Clarke, R. A.: Hydrography of the Labrador
Sea during active convection, J. Phys. Oceanogr., 32, 428–457, 2002.Pickart, R. S., Straneo, F., and Moore, G. W. K.: Is Labrador Sea Water
formed in the Irminger basin?, Deep-Sea Res. Pt. I, 50, 23–52,
10.1016/S0967-0637(02)00134-6, 2003.Piron, A., Thierry, V., Mercier, H., and Caniaux, G.: Argo float
observations of basin-scale deep convection in the Irminger sea during
winter 2011–2012, Deep-Sea Res. Pt. I, 109, 76–90, 10.1016/j.dsr.2015.12.012,
2016.Piron, A., Thierry, V., Mercier, H., and Caniaux, G.: Gyre-scale deep
convection in the subpolar North Atlantic Ocean during winter 2014–2015,
Geophys. Res. Lett., 44, 1439–1447, 10.1002/2016GL071895, 2017.Rhein, M., Steinfeldt, R., Kieke, D., Stendardo, I., and Yashayaev,
I.: Ventilation variability of Labrador Sea Water and its impact on oxygen
and anthropogenic carbon: a review, Philos. T. Roy. Soc. A, 375, 20160321,
10.1098/rsta.2016.0321, 2017.Schmidt, S. and Send, U.: Origin and Composition of Seasonal Labrador Sea
Freshwater, J. Phys. Oceanogr., 37, 1445–1454, 10.1175/JPO3065.1, 2007.
Straneo, F, Pickart, R. S., and Lavender, K.: Spreading of Labrador sea water: an
advective-diffusive study based on Lagrangian data, Deep-Sea Res. Pt.
I, 50, 701–719, 2003.Swingedouw, D., Rodehacke, C. B., Behrens, E., Menary, M., Olsen, S. M.,
and Gao, Y.: Decadal fingerprints of freshwater discharge around Greenland
in a multi-model ensemble, Clim, Dynam., 41, 695–720,
10.1007/s00382-012-1479-9, 2013.Våge, K., R. S. Pickart, V. Thierry, G. Reverdin, C. M. Lee, B. Petrie, T. A. Agnew, A. Wong, and M. H. Ribergaard: Surprising return of deep
convection to the subpolar North Atlantic Ocean in winter 2007–2008, Nat. Geosci., 2, 67–72, 10.1038/ngeo382, 2009.
Yang, Q., Dixon, T. H., Myers, P. G., Bonin, J., Chambers, D., and Van Den
Broeke, M. R.: Recent increases in Arctic freshwater flux affects Labrador
Sea convection and Atlantic overturning circulation, Nat. Commun., 7, 1–7,
10.1038/ncomms10525, 2016.Yashayaev, I. and Clarke, A.: Evolution of North Atlantic Water Masses
Inferred From Labrador Sea Salinity Series, Oceanography, 21, 30–45,
10.5670/oceanog.2008.65, 2008.Yashayaev, I. and Loder, J. W.: Recurrent replenishment of Labrador Sea
Water and associtated decadal-scale variability, J. Geophys. Res.-Oceans, 121, 8095–8114,
10.1002/2016JC012046, 2016.Yashayaev, I. and Loder, J. W.: Further intensification of deep convection
in the Labrador Sea in 2016, Geophys. Res. Lett., 44, 1429–1438,
10.1002/2016GL071668, 2017.Yashayaev, I., Bersch, M., and van Aken, H. M.: Spreading of the Labrador
Sea Water to the Irminger and Iceland basins, Geophys. Res. Lett., 34, 1–8,
10.1029/2006GL028999, 2007.