We present 2500 years of reconstructed bottom water temperatures (BWT) using
a fjord sediment archive from the north-east Atlantic region. The BWT
represent winter conditions due to the fjord hydrography and the associated
timing and frequency of bottom water renewals. The study is based on a ca.
8 m long sediment core from Gullmar Fjord (Sweden), which was dated by
210Pb and AMS 14C and analysed for stable oxygen isotopes
(δ18O) measured on shallow infaunal benthic foraminiferal species
Cassidulina laevigata d'Orbigny 1826. The BWT, calculated using the
palaeotemperature equation from McCorkle et al. (1997), range between 2.7 and
7.8 ∘C and are within the annual temperature variability that has
been instrumentally recorded in the deep fjord basin since the 1890s. The
record demonstrates a warming during the Roman Warm Period (∼350 BCE–450 CE), variable BWT during the Dark Ages (∼450–850 CE),
positive BWT anomalies during the Viking Age/Medieval Climate Anomaly (∼850–1350 CE) and a long-term cooling with distinct multidecadal
variability during the Little Ice Age (∼1350–1850 CE). The fjord BWT
record also picks up the contemporary warming of the 20th century (presented
here until 1996), which does not stand out in the 2500-year perspective and
is of the same magnitude as the Roman Warm Period and the Medieval Climate
Anomaly.
Introduction
The climate variability over last two millennia has been widely recognized as
being crucial for understanding the present and future climate responses to
anthropogenic forcing (e.g. Cunningham et al., 2013; PAGES 2K, 2013; McGregor
et al., 2015; Abram et al., 2016). To evaluate how significant regional
climate changes are or if observed temperature anomalies are unprecedented in
view of long-term climate evolution, there is a need for long historical
instrumental climate records. A major challenge for the reconstructions of
past climate changes, both using proxy data and utilizing paleoclimate modelling, is
often a lack of such long instrumental records, which if available seldom
reach beyond the 20th century. The North Atlantic region plays a paramount role
in climate variability and the global carbon budget in this
respect by modulating the Atlantic meridional overturning circulation (AMOC) (e.g.
Eiríksson et al., 2006; Lund et al., 2006; Park and Latif, 2008; Trouet
et al., 2009). The upper northern limb of the AMOC, the North Atlantic
Current (Fig. 1a), delivers heat, salt and nutrients from the tropics to the
mid- and high- latitudes and carries major parts of the volume flux into the
Nordic Seas (Hansen and Østerhus, 2000). The AMOC is thought to be linked
to the Atlantic multidecadal oscillation (AMO; Enfield et al., 2001) through
sea surface temperature variability and is connected to the decadal
variability of the North Atlantic Oscillation, (NAO; Curry and McCartney,
2001), where the NAO index is defined as the normalized sea level pressure
difference between the Icelandic low and the Azores high (Hurrell et al.,
1995). The variability of the AMOC also contributes to a multidecadal modulation
of El Niño–Southern Oscillation (ENSO) (Ortega et al., 2012 and
references therein). In addition, the North Atlantic Current passes between
the subpolar and subtropical gyres (Fig. 1a), from which it draws water;
hence, the current depends on variability occurring within both gyres (Hansen and
Østerhus, 2000). Variability of ocean temperature in the high latitude North
Atlantic and Nordic Seas are reflected in the north-western European climate and in the winter
Arctic sea ice extent (Årthun et al., 2017). Model projections predict
that the AMOC will slowdown in response to future warming and enhanced Arctic
freshwater fluxes (e.g. Schmittner et al., 2005; Ortega et al., 2012; Caesar
et al., 2018) with potential impacts on the climate, ecosystems, agriculture and
the economies of many European countries (e.g. Kuhlbrodt et al., 2009; Jackson et
al., 2015; Knox et al., 2016). Hence, high-resolution paleoceanographic
records, which preferably overlap with instrumental observations and
historical data, are needed from the eastern North Atlantic region in order
to document climate variability related to the physical properties of the North
Atlantic Current and the AMOC strength. At the same time many of the marine
records available from the region to date tend to have a low temporal
resolution due to their location in the deep sea or within the open shelf
areas. Sediment archives of temperate fjord inlets located within the eastern
North Atlantic region offer the potential for high-resolution records of the
maritime climate, because they act as sediment traps resolving climate
variability at an almost annual resolution (Howe et al., 2010). Yet to date,
there are relatively few such high-resolution paleoclimate records from the
eastern North Atlantic fjords spanning the late Holocene (e.g.
Mikalsen et al., 2001;
Klitgaard-Kristensen et al., 2004; Cage and Austin, 2010; Filipsson and
Nordberg, 2010; Hald et al., 2011; Kjennbakken et al., 2011; Faust et al.,
2016).
Map of the study area including the location of Gullmar Fjord (GF)
and the sampling site for the Ga113-2Aa and 9004 records (star) within the North Atlantic
(a) and North Sea – Skagerrak regions (b). Locations of the other proxy
records discussed in the text are shown by white circles, while some of the major ocean circulation
characteristics mentioned in the text are indicated as follows: EGC is the East
Greenland Current, NAC is the North Atlantic Current, SPG is the subpolar gyre and
STG is the subtropical gyre (a). (b) The major regional water masses and currents
are shown as follows: AW is the Atlantic Water, SJC is the South Jutland Current, NJC
is the North Jutland Current, BC is the Baltic Current and NCC is the Norwegian Coastal
Current. (c) An overview of the water column stratification in the longitudinal
profile of the Gullmar Fjord with an indication of the salinity (S) and residence
times (t) typical for each water layer (Arneborg, 2004).
Meanwhile, crucial knowledge has been gained from temperature proxy datasets
available from the North Atlantic and the Northern Hemisphere in general,
which represent either composite records of different climate characteristics
with various temporal resolution or are a combination of historical and proxy
data, and the generated data sets mostly
reflect summer conditions at higher latitudes (e.g. Moberg et al., 2005;
Gunnarsson et al., 2011; Butler et al., 2013; Cunningham et al., 2013; PAGES
2K, 2013, 2017; Sicre et al., 2014; Linderholm et al., 2015). In contrast,
based on instrumental records, increased winter temperatures have been
suggested as an important driver of the most recent warming (Cage and Austin,
2010); therefore, climate proxies incorporating winter signal are needed.
Herein, we present a bottom water temperature proxy record from the Gullmar
Fjord, on the west coast of Sweden, which illustrates the climate development
in north-western Europe over the last ∼2500 years. Strong advantages of
the presented record in this study are its high temporal (annual to
sub-decadal) resolution and the fact that a winter temperature signal is
recorded in fjord foraminiferal shells (tests) due to specific hydrographic
conditions. The reconstructed temperatures are based on stable oxygen
isotopes (δ18O) measured in tests of a shallow infaunal
foraminifer Cassidulina laevigata d'Orbigny 1826 and reflect the
deep-water temperatures in the fjord basin. The fjord has a
> 100 year long record of instrumental observations from the
deepest basin, performed since 1869 (Fig. 2a–c); furthermore, a
> 100 year long time series of air temperature observations are
also available for Stockholm, Sweden and central England. These instrumental
observations of bottom water temperatures and air temperatures are used to
evaluate the accuracy of the reconstructed climate variability for the last
century provided by the fjord sediment archive.
Study area
Gullmar Fjord is a Skagerrak fjord inlet, which is 28 km long, 1–2 km
wide and oriented south-west to north-east (Fig. 1). The maximum basin depth
is 118.6 m. The fjord is located at a critical latitude, picking up
fluctuations between cold and temperate climates, and has almost no tidal
activity. The adjacent Skagerrak largely determines the local hydrography so
that the deep (basin) water, which is typically exchanged in the fjord during
the winter, originates from the North Sea surface water flowing into the
Skagerrak with the present-day current circulation system (Svansson, 1975;
Nordberg, 1991). The 42 m deep sill at the fjord entrance restricts the
water exchange and results in a water column stratified due to salinity
differences (Fig. 1c). At the surface (< 1 m) there is a thin layer
of river water from the Örekilsälven (Fig. 1), which does not
significantly impact the fjord hydrography (Arneborg, 2004). Below, at a water depth of 1–15 m, there is a
brackish water mass (salinity (S)=24–27), primarily derived from the Baltic
current flowing northward along the Swedish west coast. The brackish water
mass has a residence time of 20–38 days in Gullmar Fjord (Arneborg et al.,
2004). A more saline water mass (S=32–33) at ∼15–50 m is derived
from the Skagerrak and has mean residence time of 29–62 days (Arneborg et
al., 2004). The last and deepest layer (> 50 m), referred to herein
as deep water or basin water, is more stagnant, with little seasonal or
inter-annual change in salinity (ranging between 34 and 35) and an inter-annual
temperature variability of 4–8 ∘C (Fig. 2a, b). The deep water
temperatures vary between years depending on the temperature of the
inflowing water mass but they remain stable seasonally (Fig. 2d). The deep-water
salinities do not vary much on a seasonal basis from the average value of 34.5
(Fig. 2b). The stratification of the water column is further strengthened
during the summer by the development of a strong thermocline, which impedes
deep-water exchange. The deep-water exchange of the fjord basin water takes
place once a year during winter, mostly between January and March, which is
determined using long-term instrumental records from the fjord (Arneborg
et al., 2004). Due to the presence of a sill, isolating the fjord deep water
from the adjacent sea, and the comparably large basin volume, the winter
temperature and salinity of the inflowing North Sea/Skagerrak water, are
“annually preserved” in the fjord basin until the next deep-water turnover
the following year (Arneborg et al., 2004). This results in a deep-water
environment characterized by winter temperatures. The benthic foraminifers
reproduce and grow in the fjord during the spring and summer (Gustafsson and
Nordberg, 2001); thus, they incorporate this annually preserved winter
temperature signal of the ambient deep water into their shells. This results
in a stable oxygen isotope signal mainly reflecting the winter temperatures of
the North Sea surface water and the Skagerrak intermediate water flowing into
the fjord during deep-water exchange.
Hydrographic measurements from Alsbäck Deep, Gullmar Fjord taken
during the 1890–2000 period below a water depth of 110 m: bottom water
temperatures (BWT) (a), salinity (b) and dissolved oxygen (c). A
snapshot of the hydrographic changes in BWT (d), salinity (e)
and oxygen (f) associated with basin water exchanges between 1992
and 1993 showing annual variability of these parameters.
The deep-water exchange in the fjord is driven by wind forcing, and largely
depends on wind direction and wind strength (Björk and Nordberg, 2003).
The latter two properties, in turn, are governed by the NAO, which is the
dominant mode of climate variability in the region during the winter
(Hurrell, 1995). In Gullmar Fjord, the higher frequency and duration of
north-easterly winds, common during the negative NAO index periods, result in Ekman
transport of surface water from the coast and facilitate coastal upwelling,
which causes the deep-water exchange (Björk and Nordberg, 2003). In
contrast, a positive NAO index causes prevailing westerly winds, which limit
the chances of deep-water renewals occurring. From the late 1970s the NAO
has been in its prolonged positive phase and is believed to be one of the
triggers of severe seasonal hypoxia (< 1 mL O2-1)
in the deep fjord basin (Nordberg et al., 2000; Björk and Nordberg, 2003;
Filipsson and Nordberg, 2004a).
After an extensive deep-water exchange event in the fjord the oxygen level
starts to decline in June, and the lowest oxygen levels normally develop
between November and January, indicating hypoxic conditions
(< 2 mL O2 L-1); however, so far anoxia has not been
recorded (Fig. 2f). The first ever documented severe hypoxic event was noted
in February 1890 by Pettersson and Ekman (1891). In the following, severe
hypoxic events were measured in 1906, 1961/1962 and 1973/1974 (Fig. 2c),
although due to the low observation frequency and duration of these events they are not
well documented. Since 1979, multiple episodes of more frequent severe
hypoxia lasting for at least 3 months have been observed. These events
occurred in 1979/1980, 1983/1984, 1987/1988, 1988/1989, 1990/1991, 1994/1995,
1996–1998, 2008, 2014/2015 and 2016 (e.g. Filipsson and Nordberg, 2004a;
Polovodova Asteman and Nordberg, 2013; SMHI SHARK-database, 2017; Nordberg et
al., unpublished data).
The severe hypoxia makes the fjord basin hostile for large burrowing
organisms but allows benthic meiofaunal communities to thrive. This lowers sediment
bioturbation and results in a well-preserved environmental sediment archive.
The fjord basin has high sediment accumulation rates, which provide a high
temporal resolution corresponding to 1–6 years per 1 cm thick sediment
sample. Finally, the fjord sediment archive is characterized by the diverse
and abundant foraminiferal fauna and dinoflagellate cysts, which have
already provided insight into climate evolution and the associated
environmental changes on the Swedish west coast during the last two millennia
(Filipsson and Nordberg, 2004a, 2010; Harland et al., 2006, 2013; Nordberg et
al., 2009; Polovodova et al., 2011; Polovodova Asteman and Nordberg, 2013;
Polovodova Asteman et al., 2013).
Age model of the studied Ga113-2Aa and 9004 records (a) and
the comparison of foraminiferal and isotopic data with core G113-091, taken at
the same location in 2009, to prove the absence of a gap between GA113-2Aa
and 9004 (b), according to Polovodova Asteman et al. (2013).
Material and methods
This study is based on a composite record of two sediment cores: GA113-2Aa
and 9004. Both cores were collected at a water depth of 116 m at the same
site in the deepest Gullmar Fjord basin (58∘17.570′ N,
11∘23.060′ E) (Fig. 1) for which the long-term hydrographic
observations are available (Fig. 2a–c). Core 9004 (731 cm long) was taken
with a gravity corer (∅=7.6 cm) onboard R/V Svanic in
July 1990. Core GA113-2Aa (60 cm long), which had an intact sediment–bottom
water interface, was recovered using a Gemini corer (∅=8 cm) in
June 1999 from the R/V Skagerak. In the laboratory both cores were
split into two halves and sectioned at 1 cm intervals. One half was used for
bulk sediment geochemistry (total carbon – TC, total nitrogen – TN and carbon to nitrogen ratio – C / N), stable oxygen and carbon isotopes, dinoflagellate cysts and benthic
foraminiferal faunal analyses. Another half was stored as an archive at the
Department of Geosciences, University of Gothenburg. The TC and stable carbon
isotope data from both cores are published in Filipsson and Nordberg (2010),
dinoflagellate cysts data are discussed in Harland et al. (2006, 2013), while
C / N and foraminiferal assemblage data are presented in Filipsson and
Nordberg (2004a), Polovodova et al. (2011) and Polovodova Asteman et
al. (2013). We also present data from the gravity core G113-091, collected at
the same location as GA113-2Aa and 9004 onboard R/V Skagerak in
September 2009; this data is only used herein (similar to our previous study)
to create a composite age model for the GA113-2Aa and 9004 cores (Polovodova
Asteman et al., 2013; see below).
Stations with collected sediment core tops and δ18O
analyses on living (Rose Bengal stained) Cassidulina laevigata.
In addition to the above-mentioned cores, we also use six surface samples
(0–1 cm) collected at five stations in the Skagerrak (OS4, OS6, OS14, 9202
and 9205) and one station in the Gullmar Fjord (G113-091a: the same location
as for GA113-2Aa and 9004) in 1992–93 and 2009, respectively (Fig. 1b, c;
Table 1). All surface samples were stained using Rose Bengal to distinguish
individuals presumably living at the moment of sampling from the empty
foraminiferal shells.
AMS 14C dates obtained for the gravity core 9004 and calibrated
calendar ages. All dates presented in Filipsson and Nordberg (2010) and
Polovodova Asteman et al. (2013) were re-calibrated using Calib 7.10 (Stuiver
et al., 2017), the Marine13 calibration dataset (Reimer et al., 2013), and
ΔR=100±50.
* Dates not
used in the final age model due to age reversals.
Sediment core dating and age model
The age model for the composite GA113-2Aa – 9004 record was previously
published in Filipsson and Nordberg (2010) with further revisions by
Polovodova et al. (2011) and Polovodova Asteman et al. (2013). Eleven intact
marine bivalve shells were recovered in life position from core 9004 and were
subject to AMS 14C analysis (Fig. 3a; Table 2). All 14C dates were
obtained through analysis at the Ångström Laboratory (Uppsala
University, Sweden) and were originally calibrated using the marine
calibration curve (Reimer et al., 2004; Bronk Ramsey, 2005). Ages were
normalized to δ13C of - 25 ‰ according to Stuiver and
Polach (1977), and a correction corresponding to δ13C = 0 ‰ (not measured) versus Pee Dee Belemnite (PDB) has
been applied. Herein we present ages recalibrated using Calib Radiocarbon
Calibration software v. 7.1 (Stuiver et al., 2017:
http://calib.org/calib/; last access: 15 March 2017), the most recent
marine calibration curve (Reimer et al, 2013) and a reservoir age of
500 years (ΔR=100±50),
which was obtained from pre-bomb marine bivalve shells from the Gullmar
Fjord, provided by the natural history museums in Gothenburg and Stockholm
(Nordberg and Posnert, unpublished data). All ages are presented as median
probability with a 1-σ error margin (Table 2). Two dates at 98 and
313 cm showed minor age reversals and were omitted from the final age model
(Table 2). The GA113-2Aa core was dated using 210Pb and a constant rate
of supply (CRS) model (Appleby and Oldfield, 1978), which suggested that the
core material was deposited between ca. 1915 and 1999 (Fig. 3a). For details
regarding the GA113-2Aa age model see Filipsson and Nordberg (2004a).
Together, cores GA113-2Aa and 9004 proved to be a continuous sediment
record with no gap in between based on the correlation of the stable carbon
isotopes (δ13C) and benthic foraminiferal species C. laevigata, Adercotryma glomerata (Brady, 1878) and Hyalinea balthica (Schröter in Gmelin,
1791) with respective data from core G113-091 (Fig. 3b herein; Polovodova
Asteman et al., 2013; Polovodova Asteman and Nordberg, 2013). The composite
record of GA113-2Aa and 9004 spans from approximately 350 BCE to 1999 CE
(Table 2, Fig. 3a), and includes the late Holocene climate events such as the
Roman Warm Period (RWP: ∼350 BCE–450 CE), the Dark Ages Cold Period
(DA: ∼450–850 CE), the Viking Age/Medieval Climate Anomaly (VA/MCA:
∼850–1350 CE), the Little Ice Age (LIA: ∼1350–1850 CE) and
the contemporary warming from 1850 CE to present (Lamb, 1995; Filipsson and
Nordberg, 2010; Harland et al., 2013, 2017; Polovodova Asteman et al., 2013).
We add the Viking Age to the Medieval Climate Anomaly following the approach
of Filipsson and Nordberg (2010), based on historical evidence that warming
in northern Europe began earlier than 1000 CE, which allowed Vikings to
reach the north-east coast of England and loot the monastery of Lindisfarne in
793 CE (Morris, 1985). For further details on the chronology of the GA113-2Aa and 9004 cores
see Filipsson and Nordberg (2004a), Polovodova et al.,
(2011) and Polovodova Asteman et al. (2013).
Combining the long gravity core with the 60 cm long Gemini core, which
includes the sediment–bottom water interface and the intact core top,
resulted in a high-resolution temporal record of almost 1-year cm-1
sample for the upper part of the record and < 10 years cm-1
sample for the deepest part of the record. Calculations from the 210Pb
analyses and the AMS-14C dates suggest sediment accumulation rates of
∼9 mm year-1 in the most recent sediments and approximately ∼2.8 mm year-1 in the compacted deepest part of the gravity core
(Fig. 2). Hence, due to high accumulation rates the upper 60 cm of the
record can be directly compared to instrumental hydrographic and
meteorological data (Figs. 7 and 8).
Stable oxygen isotopes
We measured δ18O on tests of the shallow infaunal foraminifer
Cassidulina laevigata from the core top samples and from the
ca. 8 m long G113-2Aa – 9004 record (Fig. 1b). Between 12 and 20 specimens
of Cassidulina laevigata were picked from each sample for the
analysis. In total 6 and 425 samples were analysed for stable oxygen isotopic
composition for the surface sediments and composite G113-2Aa – 9004 record,
respectively. All samples were measured at the Department of Geosciences,
University of Bremen, Germany, using a Finnigan Mat 251 mass spectrometer
equipped with an automatic carbonate preparation device. Isotope composition
is given in the usual δ-notation and is calibrated to Vienna Pee Dee
Belemnite (V-PDB) standard. The analytical standard deviation is
< 0.07 ‰ for δ18O based on the long-term
standard deviation of an internal standard (Solnhofen limestone).
The temperature was reconstructed using the salinity–δ18Ow relationship established by Fröhlich et al. (1988)
(Eq. 1), which is representative for this region (Filipsson, unpubl. data).
An average salinity value of 34.4 (range 33–35) was used in Equation 1,
based on instrumental measurements between 1896 and 1999 for the fjord
deep water (station Alsbäck Deep). The salinity (S) was assumed to be
constant over the investigated time period.
δ18Ow=0.272×S-8.91
To calculate temperatures the paleotemperature equation by McCorkle et
al. (1997) was applied (Eq. 2). This equation is more appropriate for the
temperature range observed in temperate fjord basin than the more commonly
used linear equation by Shackleton (1974), which produces unrealistically
high temperatures in our study (see results section). The bottom water
temperature in degrees Kelvin (T∘K) was calculated as follows:
T∘K=2.78⋅103ln(δ18Oc+10000.97006⋅δ18Ow-29.94+1000)+2.89103,
where δ18Oc stands for stable oxygen isotopic ratio
18O/16O measured in calcite tests of C. laevigata,
while δ18Ow is the isotopic composition of water
calculated from Equation 1 and converted from SMOW to V-PDB by subtracting
0.27 ‰ (Bemis et al., 1998).
Finally, to convert reconstructed temperatures to degrees Celsius, Eq. (3)
was used:
T∘C=T∘K-273.15.
Since 1990 C. laevigata has become a rare species in the Gullmar
Fjord deep basin (Fig. 6), which has resulted in a short gap in the most recent
part of the record (see discussion). Similar gaps in δ18O and,
hence, in bottom water temperature data are also seen for the earlier part of
the record and are due to the absence or very low abundances of C. laevigata (Fig. 6).
Comparison of reconstructed temperatures and δ18O values
measured in stained C. laevigata from the core tops collected in
Gullmar Fjord (G113-091) and the Skagerrak (OS4, OS6, OS14, 9202, 9205) to
hydrographic temperature data (a) and to δ18O predicted
from the palaeotemperature equation (b) by McCorkle et al. (1997).
(c) Temperature vs. δ18Oc–δ18Ow, together
with the paleotemperature equations from Shackleton (1974), Hays and
Grossman (1991), Kim and O'Neil (1997), McCorkle et al. (1997) and Bemis et al. (1998).
Hydrographical and meteorological instrumental data
Long-term hydrographical instrumental data for temperature, salinity and
dissolved oxygen concentration (O2) for the fjord basin (average
for 110–118 m water depth (w.d.)) were extracted from
the Swedish Meteorological and Hydrological Institute (SMHI) SHARK database
(https://www.smhi.se/klimatdata/oceanografi/havsmiljodata/marina-miljoovervakningsdata;
last access: 15 March 2017). Some of the Gullmar instrumental data is also
available from the Water Quality Association of the Bohus Coast (BVVF)
(http://www.bvvf.se/; last access: 15 March 2017), while the data prior
to 1958 was sourced from Engström (1970). The Skagerrak hydrography data
for the stations adjacent to OS4-6, 9202, 9205 and OS14 were obtained from
the International Council for the Exploration of the Seas (ICES:
http://www.ices.dk/marine-data/; last access: 15 March 2017).
Meteorological observations of air temperature were also obtained for
Stockholm (https://www.smhi.se/klimatdata; last access: 15 March 2017) and central England
(http://www.metoffice.gov.uk/; last access: 16 March 2017), which both have the longest historical meteorological
records going as far back as the 18th century.
ResultsCore tops
To obtain an error estimate and to facilitate the choice of the
paleotemperature equation we used living (stained) tests of
Cassidulina laevigata from the core top samples collected in the
Gullmar Fjord and the adjacent Skagerrak. Calculated bottom water
temperatures based on the δ18Oc values from the living
(stained) C. laevigata were compared to ICES and SMHI hydrography
data from the adjacent stations (Fig. 4a). The δ18Oc
values predicted from the chosen equation (see below) were also used to
estimate the reliability of our temperature reconstruction (Fig. 4b).
Cassidulina laevigata has previously been suggested to calcify
0.19 ‰ lower than equilibrium (Poole et al., 1994). Our
δ18Oc data from the core tops demonstrate an offset,
ranging between 0.01 ‰ and 0.27 ‰ (mean 0.15 ‰),
compared with δ18Oc predicted using the
palaeotemperature equation from McCorkle et al. (1997) (Fig. 4b). Applying
the mean correction of +0.15 ‰ to the Gullmar
δ18Oc record results in bottom water temperatures ∼0.5–1 ∘C higher than those recorded by instrumental observations
in the fjord (Fig. 2a), while uncorrected δ18Oc values
produce temperatures close to observations. Taking the latter into the
account and because, based on available data, it is difficult to estimate how
large the correction should be, we further report the uncorrected
δ18Oc values for both the core tops and the sediment
cores. Furthermore, we use a median value (0.7 ∘C) of the range of
the produced temperature offset (Fig. 4a) as an error margin for our
paleotemperature reconstructions (Figs. 5–6).
A 2500-year long δ18O record (a) and reconstructed
winter bottom water temperatures, BWT (b) from Gullmar Fjord. Thick lines
show the 3-point running mean for both curves, and dashed lines indicate (a) a
long-term average of 2.4 ‰ for δ18O record
and (b) 5.4 ∘C – a mean for instrumental bottom water temperatures
registered between 1961 and 1990. Grey shaded areas in BWT indicate a median
offset (0.7 ∘C) in instrumental versus reconstructed temperatures
obtained for Rose Bengal stained C. laevigata from the core tops (see Fig. 4a), used
herein as an error margin. (c) The box and whisker plot shows a range for
instrumental BWT observations performed from 1890 to 1999 and measured at
more regular intervals from the 1960s, the data is from water depths ≥110 m in the fjord's deepest basin (Alsbäck Deep). The middle, the upper
and the lower horizontal lines in the box indicate the median value and 75th and 25th percentiles,
respectively. Abbreviations are as follows: RWP represents the Roman
Warm Period, DA represents the Dark Ages, VA/MCA represents the Viking Age/Medieval Climate
Anomaly and LIA represents the Little Ice Age.
Instrumental temperature data from ICES and SMHI were used to calculate
δ18Oc– δ18Ow for the core top
samples to facilitate the choice of a paleotemperature equation. Plotting
δ18Oc–δ18Owversus observed
temperature data for different paleotemperature equations (Fig. 4c) allowed
for an estimation of which of the equations gives the best possible agreement with the
core top data; hence, the most appropriate equation could be chosen for temperature
reconstructions. Figure 4c shows that δ18O values from north-western
Skagerrak (OS4 and OS6) are clearly in better agreement with equations by
Hays and Grossman (1991) and McCorkle et al. (1997), while the central
Skagerrak samples (9202 and 9205) plot close to the linear equation by
Shackleton (1974). The samples from Gullmar Fjord (G113-091) and the OS14
station, collected just outside the fjord, occupy a space in between the
Shackleton equation and those by Hays and Grossman (1991) and McCorkle et
al. (1997). This suggests that applying the Shackleton equation for Gullmar
Fjord and Skagerrak will result in temperatures higher than observations,
which has been also observed for Cibicidoides and Planulina
from the Florida Straits (Marchitto et al., 2014). Indeed, when testing the
Shackleton equation on our dataset, the temperatures are warmer than the ICES
hydrographic observation data by 1.5–2 ∘C. In contrast, the
equation by Bemis et al. (1998) applied to the core top
δ18Oc data produces the coldest temperatures, which are
0.9–1.9 ∘C colder than observations. In turn, it appears that by
using Hays and Grossman (1991) or McCorkle et al. (1997) equations, the
corresponding calculated temperatures come closer to observations. Both
equations are nearly identical for the temperature range from 3 to 8 ∘C
(Fig. 4c) observed between 1890 and 2001 (Fig. 2), and by exercising both
equations on the Gullmar Fjord δ18Oc record almost
identical paleotemperature curves are produced. This is rather curious since
the equation from Hays and Grossman (1991) is based on meteoric calcite of
non-biogenic origin. For this reason, in the current study we apply the
McCorkle et al. (1997) equation for the paleotemperature reconstructions.
Reconstructed bottom water temperatures (BWT) shown as an anomaly
against the 1961–1990 instrumental mean of 5.4 ∘C from Gullmar
Fjord and compared against other temperature proxy records: annual Northern
Hemisphere temperatures (Moberg et al., 2005), bottom water temperatures
from Malangen Fjord in north-western Norway (Hald et al., 2011) and Loch Sunart in
Scotland (Cage and Austin, 2010), spring sea surface temperatures from
Chesapeake Bay, eastern North Atlantic Ocean (Cronin et al., 2003), annual
temperatures reconstructed for continental Europe (PAGES 2K, 2013) and the
reconstructed NAO record from Trondheim Fjord, western Norway (Faust et al.,
2016). Also shown, are relative abundances of foraminifer Cassidulina laevigata in the fjord with
abundance minima and respective gaps in temperature reconstruction linked to
the positive NAO index (arrows). For the locations of these proxy records see
Fig. 1a and for abbreviations see the caption of Fig. 5. Grey shaded areas in
the Gullmar Fjord BWT anomalies indicate a median offset (0.7 ∘C) in
instrumental versus reconstructed temperatures (see Fig. 4a) obtained for
Rose Bengal stained C. laevigata from the core tops, used herein as an error margin.
Blue and pink boxes depict a short-lived cooling at
∼1250 CE and a warm interval between ∼1570 and 1700 CE,
respectively (as discussed in the text).
Composite record of the G113-2Aa and 9004 sediment cores
The δ18O record from the Gullmar Fjord shows both decadal and
centennial variability for the last 2500 years (Fig. 5a) and can be divided
into five major isotopic intervals.
For the lower part of the record at
802–592 cm, corresponding to ∼350 BCE–450 CE, the δ18Oc values are generally lower (∼2.4 ‰) than
the long-term average of 2.7 ‰.
Between 592 and 475 cm (∼450–900 CE) the δ18O record demonstrates a considerable
variability (Fig. 5a), starting with higher δ18Oc
(2.8 ‰–3 ‰) at 592–574 cm (∼450–525 CE), which
then become lower (∼2.4 ‰) at 574–529 cm (∼525–700 CE) and increase again (∼3.0 ‰) between 529 and
497 cm (∼700–825 CE).
The 475–302 cm interval (∼900–1350 CE) once again displays lower δ18Oc (∼2.4 ‰–2.5 ‰), which are below the long-term average.
From 302 to 53.5 cm (∼1350–1900 CE) the stable oxygen isotope
record increases again with the majority of the δ18Oc
values being ∼3.1 ‰–3.2 ‰ and exceeding the
long-term average. Within this interval the highest δ18O values of
> 3.2 ‰ are found between 300 and 170 cm (∼1350 CE–1580 CE).
Finally, the δ18O record becomes lower
again (∼2.4 ‰) between 53.5 and 5 cm (∼1900 and
1996 CE). The δ18O data for the samples between 5 and 0 cm
(1996–1999) are missing because we did not find enough
Cassidulina laevigata specimens to perform isotopic analyses.
Shifts of ∼0.25 ‰ in δ18Oc occur
throughout the Gullmar Fjord δ18O record, which according to the
equation from McCorkle et al. (1997) (used herein) may potentially indicate a
temperature variability of ∼1∘C. The corresponding salinity
change is rather small (0.02), and was calculated using the mixing line by
Fröhlich et al. (1988) and by applying the δ18Oc
range of 2.6–2.85 and a corresponding temperature range of
4.9–5.9 ∘C. Such salinity changes are well within the amplitude of
inter-annual variability (1–1.5), recorded by instrumental salinity
observations since the 1890 (Fig. 2b). Foraminifera precipitate their tests
during several months (e.g. Filipsson et al., 2004); thus, they integrate the
inter-monthly salinity signal, which together with annual variability is
minimal according to the instrumental data. For the upper part of the record,
the 1 cm sediment slice integrates one or possibly two growing seasons of
C. laevigata and subsequently records a potentially higher variability of
both salinity and temperature. However, in the deepest part of the record a
single 1 cm sample may correspond to ∼7–10 years and is more
likely to average inter-annual salinity and temperature variability providing
“a more smoothed” signal.
Stable carbon isotope (δ13C) data from the composite G113-2Aa –
9004 record (Filipsson and Nordberg, 2010) were plotted against the oxygen
isotope data presented herein, to investigate the potential relationship
between the two, e.g., due to different water masses (Supplement Fig. S1). No
such relationship was found (Fig. S1), which indicates that our
δ18O record mainly reflects fjord deep-water temperatures.
Comparison of the winter bottom water temperatures (BWT)
reconstructed from the Gullmar Fjord record to instrumental basin water
temperatures measured in the deepest fjord basin: the annual mean
(a), the mean for May–August (b) and the mean for
January–March (c). All instrumental BWT curves (black) are shown as 10-point running means (10-p. r.m.).
Reconstructed bottom water temperatures (BWT)
The resulting calculated bottom water temperature record is plotted as both
absolute temperature values (Fig. 5b) and as an anomaly from the mean value
(5.4 ∘C), based on the instrumental temperatures observed between
1961 and 1999 (Fig. 6). With very few outliers, the reconstructed temperature
range (2.7–7.8 ∘C) is within the present-day annual variability,
documented from instrumental temperature measurements in the fjord's deepest
basin since 1890 (Figs. 2a–c, 5c). To further prove that our record
represents a winter signal rather than summer conditions (as most biological
proxies) we compare the obtained BWT record to instrumental temperatures
recorded in the fjord deep water during summer and winter. When performing
such a comparison, instead of the commonly used June–August (JJA)
temperatures for designation of meteorological summer, one has to consider
the observations during May–August, when foraminifera precipitate their
calcite (Gustafsson and Nordberg, 2001; Filipsson et al., 2004), which is
used herein for stable isotope analysis. Likewise, instead of using the months typically employed
for the definition of meteorological winter (December–February: DJF), when comparing
our record to instrumental data we use January–March. January–March defines
“hydrographic winter” in the fjord, and is associated with months when
deep-water exchanges occur (see Section 2 of this paper).
Observed annual temperatures registered between 1890 and 1996 (which
corresponds to the uppermost part of the composite G113-2Aa – 9004 record)
vary between 3.0 and 8.3 ∘C, which gives an amplitude of
5.3 ∘C. Corresponding instrumental 1890–1996 temperatures for the
foraminiferal growth season in the fjord (May-August: see above) show a
4.1–7.2 ∘C range with an amplitude of 5.4 ∘C. When
studying the reconstructed temperatures over the last 2500 years the
corresponding amplitude, i.e., the difference between the maximum
(7.8 ∘C) and the minimum (2.7 ∘C) temperatures is
5.1 ∘C (Fig. 5b). Furthermore, when plotting the reconstructed bottom water
temperatures for the 1890–1996 period versus corresponding instrumental
bottom water temperatures as annual averages and means for May–August
(Fig. 7b) and January–March (Fig. 7c), the calculated bottom water
temperatures and hydrographic data agree with each other rather well in terms
of amplitude. However, an increased agreement is reached when comparing the
reconstructed data to the hydrographic winter (January–March) temperature
(Fig. 7c), which is not surprising considering the fjord hydrography and a
season when deep-water exchanges typically occur (see Section 2).
Hence, the Gullmar Fjord δ18O-based temperature record reflects the
winter temperature variability of surface water in the North Sea.
From the reconstructed Gullmar Fjord temperature record five bottom water
temperature intervals can be recognized (Figs. 5b, 6), in parallel to the
isotopic intervals mentioned above.
From ∼350 BCE to 450 CE the
fjord bottom water temperatures are consistently above 5.4 ∘C, the
1961–1990 mean.
Between 450 CE and 850 CE the record fluctuates
between positive temperature anomalies (∼450–650 CE) and negative
anomalies (∼650–850 CE) reaching minimum value at ∼750 CE.
At ∼850–1300 CE the bottom water temperatures are again above
average with a short negative anomaly around 1200–1250 CE.
The period
between ∼1300 CE and 1850 CE in the Gullmar Fjord record is
unprecedentedly cold for the last ∼2500 years with the majority of
temperature anomalies being negative and reaching the minimum value around
∼1350 CE (Fig. 6).
Finally, from ∼1850 CE towards present
day the record is characterized by consistently positive bottom water
temperature anomalies, which are comparable in the amplitude to the high
anomalies found at ∼350 BCE–450 CE.
Gaps in the record due to absent/rare Cassidulina laevigata
Some intervals in the G113-2Aa – 9004 record were barren of C. laevigata tests; hence, for these intervals, δ18O values and the
corresponding reconstructed bottom water temperature data are missing. These
intervals are ∼130–120 BCE, ∼725–740 CE, ∼1260–1265 CE, ∼1273–1277 CE, ∼1340 CE and ∼1996–1999
(Fig. 6). The most recent period of absent/rare C. laevigata in the
Gullmar Fjord coincides with higher bottom water temperatures and frequently
occurring severe hypoxia (see introduction and discussion), as registered by
the instrumental measurements in the fjord's deepest basin (Fig. 2c).
Discussion
The Gullmar Fjord winter bottom water temperature record shows both
centennial and multidecadal variability and has a striking resemblance to
climate periods (see below) historically known in northern Europe over
the last 2500 years (e.g. Lamb, 1995; Stuiver et al., 1995; Moberg et al.,
2005; Filipsson and Nordberg, 2010; Helama et al., 2017). The record
demonstrates periods of temperature variability, which correspond to the
Roman Warm Period (∼350 BCE–450 CE), the Dark Ages cold period
(∼450–850 CE), the warm Viking Age/Medieval Climate Anomaly (∼850–1350 CE), the colder Little Ice Age (∼1350–1850 CE) as well as
the warmer conditions during the 20th century (∼1850 CE–present).
There is an overall cooling trend in the Gullmar Fjord temperature record for
the last 2500 years, which is consistent with other climate proxy records for
this period (e.g. Lebreiro et al., 2006; Eiriksson et al., 2006; Hald et al.,
2011; McGregor et al., 2015). Among forcing mechanisms for the late Holocene
climate variability in the North Atlantic region changes in temperature and
the influx of the Atlantic Water to the region (e.g. Nordberg, 1991; Hass, 1996;
Klitgaard-Kristensen et al., 2004; Eiriksson et al., 2006; Lund et al.,
2006), radiative forcing (Jiang et al., 2005; Hald et al., 2007), volcanic
activity (Otterå et al., 2010; McGregor et al., 2015), land use changes
and increased greenhouse gas emissions (e.g. Masson-Delmotte et al., 2013;
Abram et al., 2016) are suggested. In addition, there is a strong coupling
between atmospheric and ocean circulation, which is linked to the NAO variability.
The NAO influences the strength and frequency of moist westerly winds that bring
precipitation to northern Europe and has even been suggested to induce
multidecadal-scale changes in the AMOC (Dickson et al., 1996), which on
centennial scales are linked to the late Holocene major climate extremes
(Bianchi and McCave, 1999). Below we discuss each of the climate extremes in
detail and compare our record to available temperature proxy data from other
settings, highly influenced by the multidecadal NAO variability and climate
changes associated with it.
The Roman Warm Period (prior to ∼450 CE)
The fjord record shows consistently positive bottom water temperature
anomalies during the Roman Warm Period (RWP) when compared to
5.4 ∘C, the annual mean for 1961–1999 (Fig. 6). The RWP is often
associated with increasingly warm and dry summers both on the British Isles
and in central Europe and is linked to the expansion of the Roman Empire
(Lamb, 1995; Wang et al., 2012). The RWP warming coincided with a more
vigorous flow of the Iceland Scotland Overflow Water, which is an important
component of the AMOC (Bianchi and McCave, 1999). Other studies report an increase in the
contribution of the Atlantic water to the East Greenland shelf, a reduced sea
ice concentration and an increase in the export of fresh water from the
Arctic with the East Greenland Current (Fig. 1a), which are all thought to be
linked to a shift from the negative to a positive NAO after ∼500 BCE/0 CE and changes in the AMO regime (e.g. Perner et al., 2015 and
references therein; Kolling et al., 2017). Harland et al. (2013) analysed
dinoflagellate cysts from the same composite core as presented herein and,
based on observed changes in species composition, suggested that sea surface
temperatures (SSTs) in the fjord were > 10 ∘C during the
RWP, as compared to present-day SSTs of ∼9∘C (SMHI, 2017).
Other studies suggest SSTs of 6–10 ∘C for the waters off northern
Iceland (Sicre et al., 2011), 10.7–12.6 ∘C for the Vøring
Plateau, Norwegian Sea (Risebrobakken et al., 2011),
> 13 ∘C off north-western Scotland (Wang et al., 2012)
and > 15 ∘C for the Rockall Trough, north-eastern
Atlantic (Richter et al., 2009) during this period. Furthermore, for the
coastal north-western Atlantic (Chesapeake Bay) SSTs as high as
12–15 ∘C have been reported (Cronin et al., 2003).
For the adjacent Skagerrak an increase in both intermediate and bottom water
temperatures is reported based on Mg / Ca data on benthic foraminiferal
species Melonis barleeanus (Butruille et al., 2017). Butruille et al. (2017) demonstrate a ∼2∘C temperature increase and report a temperature range of ∼6–8 ∘C during the RWP. In a 2000-year long temperature record from
the Malangen Fjord, north-western Norway (Hald et al., 2011), the RWP is
characterized as “a warm period with stable bottom water temperatures”. The
Malangen Fjord record is based on δ18O measured on
Cassidulinaneoteretis Seidenkrantz 1995 and documents a
bottom water temperature range of 5.5–7.5 ∘C (Hald et al., 2011).
Both the Skagerrak and Malangen Fjord studies agree well with our dataset,
which demonstrates a temperature increase of ∼2.5∘C,
resulting in a 5.4–7.9 ∘C temperature range during the RWP for the
Gullmar Fjord deep water (Fig. 5). The somewhat higher upper range limit of
the RWP bottom water temperatures in the Skagerrak and Malangen Fjord,
compared to our data, may be explained by the more direct influence of the
more temperate Atlantic water at those sites, which may be less prominent in
our study area as it is more land-locked and has a stronger continental
influence. Also given that our record reflects winter temperatures, its lower
BWT temperature range during the RWP is quite reasonable.
When comparing our data to the major temperature synthesis efforts undertaken for
the last two millennia, it becomes evident that our RWP reconstruction seem
to disagree with the Northern Hemisphere temperature record of Moberg et
al. (2005), which is mostly characterized by the negative RWP temperature
anomalies (Fig. 6). In contrast, the warming seen in the Gullmar Fjord
dataset is consistent with the PAGES 2K temperature synthesis for
continental Europe (Fig. 6), which also reports a distinct warming
corresponding to ∼2–3 ∘C temperature increase during the RWP
(PAGES 2K, 2013).
The Dark Ages Cold Period (∼450–850 CE)
Our record displays variable bottom water temperatures in the fjord during
the Dark Ages (Figs. 5–6), which is initiated by a short-lived negative
anomaly at ∼400–450 CE. This anomaly then switches to positive values (∼450–650 CE)
before once again becoming negative at ∼650–850 CE. The Dark Ages
Cold Period (DACP) is commonly linked to a large-scale human migration in
central Europe (Lamb, 1995; Büntgen et al., 2011). The DACP was
contemporaneous with a reduced flow of the Iceland Scotland Overflow Water
(Bianchi and McCave, 1999), low solar activity, low pollen influx (Desprat et
al., 2003), glacial advance (Lamb, 1995) and a negative mode of the NAO (e.g.
Seidenkrantz et al., 2007; Orme et al., 2015; Helama et al., 2017). Summer
temperatures < 10 ∘C in French Alps (Millet et al., 2009),
increased humidity in northern Europe (Barber et al., 2004) and a widespread
abandonment of arable lands and cultivation in south-western Norway
(Salvesen, 1979) were also documented for this period. Furthermore,
Seidenkrantz et al. (2007) report a warming of subsurface waters off western
Greenland during the DACP attributed to a stronger Atlantic component of the
West Greenland Current and a negative NAO.
There is also some cooling during the DACP indicated for the intermediate and
deep water in the adjacent Skagerrak (Butruille et al., 2017) but the lower
temporal resolution makes it difficult to directly compare the Skagerrak
record with ours. In contrast, variable SSTs during the Dark Ages are
reported by some North Atlantic records (Sicre et al., 2011; Risebrobakken et
al., 2011), with timing similar to the variability of the Gullmar Fjord
temperatures (see above). Variable bottom water temperatures are also
reported for the Malangen Fjord with a range (5.5–7.5 ∘C)
relatively close to our results (∼4–8 ∘C). There is also some
fluctuation between cooling and warming with a ∼3–4 ∘C
amplitude in a Mg / Ca-based SST record from Chesapeake Bay (Cronin et
al., 2003) and the DACP temperatures reconstructed for
continental Europe (PAGES 2K, 2013).
The Viking Age/Medieval Climate Anomaly (∼850–1350 CE)
After the Dark Ages the bottom water temperature anomalies in Gullmar Fjord
become positive between ∼850 CE and 1350 CE, which fits well with the
onset of the warming during the VA/MCA. The warm MCA is believed to be
associated with a positive NAO index (e.g. Trouet et al., 2009; Faust et al.,
2016), which is likely to have strengthened the AMOC (Bianchi and McCave, 1999) and
resulted in an increased transport of heat and moisture to the higher
latitudes. The MCA also coincided with grand solar maximum at 1100–1250 CE
(Zicheng and Ito, 2000) and its temperature optimum occurred between 1000 CE
and 1300 CE, when there was a sharp temperature maximum in most of Europe
(Lamb, 1995).
The mean annual northern hemispheric and continental Europe temperature
records (Moberg et al., 2005; PAGES 2K, 2013) show the onset of warming as
early as between ∼850 and 950 CE, with distinct warmth peaks reached
around 1000 and 1100 CE and the MCA termination around 1300 CE, which all
agrees rather well with our data (Fig. 6). The Malangen Fjord record also already
shows warming before 800 CE, which terminates around 1250 CE
(Hald et al., 2011), a century earlier than in the Gullmar Fjord record.
Despite the inconsistency in timing, which likely results from dating
uncertainties (which may be the case for both studies), the two fjord records
agree with each other rather well in terms of reconstructed bottom water
temperature ranges for this period: 5.4–7.6 ∘C for the Gullmar
Fjord and 5.5–7.1 ∘C for the Malangen Fjord. In the adjacent
Skagerrak both intermediate and deep-water temperatures are reported to
increase from ∼6 to 8 ∘C (Butruille et al., 2017) but sampling
resolution of the former is too low for the MCA period. In turn, bottom water
temperatures in Loch Sunart, in Scottland, increased by ∼1.2∘C during the MCA (Cage and Austin, 2010), which is also within
the abovementioned ranges. Furthermore, an increase of a similar magnitude during
the MCA is reported for the sea surface temperatures in the North Atlantic
(Cunningham et al., 2013).
An interesting feature in the Gullmar Fjord record of the VA/MCA is a
presence of a short-lived cooling centred at ∼1250 CE, before the final
peak of warmth at 1250–1350 CE (Fig. 6: see blue box). Such short cooling
during the MCA is also documented for both the eastern and western Atlantic
coasts (Chesapeake Bay: Cronin et al., 2003; Loch Sunart: Cage and Austin,
2010) but with a slightly different timing, either due to dating
uncertainties or the application of different temperature proxies (Mg / Ca
vs. δ18O).
Comparison of reconstructed winter bottom water temperatures (BWT)
from Gullmar Fjord to meteorological observations of winter air temperatures
recorded for Stockholm (stippled line) and central England (solid line
without symbols).
The Little Ice Age (∼1350–1850 CE)
From ∼1350 to ∼1850 CE our record shows winter bottom water
temperatures 2–3 ∘C lower than the instrumental annual mean for
1961–1999 (Fig. 6). Many other proxy records report cooling of a similar
magnitude or even stronger in the North Atlantic during the LIA (e.g. Stuiver
et al., 1995; Cronin et al., 2003; Klitgaard Kristensen et al., 2004;
Eiríksson et al., 2006; Hald et al., 2011; Sicre et al., 2011). The
PAGES 2K synthesis of marine palaeoclimate records spanning the past
2000 years also identified a robust global surface ocean cooling with the
coldest conditions from 1400 to 1800 CE (McGregor et al., 2015). The Little
Ice Age is commonly associated with glacial advances in the Arctic and alpine
regions (Porter, 1986; Miller et al., 2012) in response to reduced solar
activity (Mauquoy et al., 2002) and summer insolation (Wanner et al., 2011),
increased volcanism (Miller et al., 2012), negative North Atlantic
Oscillation (e.g. Trouet et al., 2009; Faust et al., 2016) and the reduced
strength of the AMOC (e.g. Bianchi and McCave, 1999; Klitgaard Kristensen et
al., 2004; Lund et al., 2006). There is also a growing evidence for a
stronger Siberian high prevailing from 1450 to 1900 CE based on increased
Na2+ content in the GISP2 record from Greenland (Mayewski et al.,
1997; Meeker and Mayewski, 2002). The onset of the LIA (∼1350 CE) on
the Swedish west coast also coincided with an outbreak of the Black
Death, which decreased the population by 50–60 % and resulted in
large-scale farm abandonment with negative implications for land use
(Harrison, 2000).
For the Gullmar Fjord a general cooling during the LIA has previously been
suggested based on increased abundances of cryophilic dinocysts (Harland et
al., 2013) and benthic foraminifer Adercotryma glomerata, which
prefer bottom water temperatures < 4 ∘C (Polovodova
Asteman et al., 2013). This agrees rather well with the data presented
in this study, which show temperatures as low as ∼3.4–4.4 ∘C around
1350, 1500, 1550 and from 1700 to 1850 CE with a general temperature range of
2.9–6.6 ∘C for the whole LIA period (Fig. 5). Based on
foraminiferal faunal and δ13C data Polovodova Asteman et al.
(2013) divided the LIA into two distinct phases in the Gullmar Fjord:
(1) 1350–1650 CE and (2) 1650–1850 CE, which were separated by a short-lived warming
centred at ∼1650 CE. The reconstructed temperatures also show a
short milder episode based on positive anomalies between ∼1570 and
1700 CE (Fig. 6: see pink box). A similarly warm (but slightly displaced in
time) event is visible in other climate records (Fig. 6) from the North
Atlantic and Northern Hemisphere (Cronin et al., 2003; Moberg et al., 2005;
Cage and Austin, 2010; Hald et al., 2011). This suggests that this short-lived
warming was a larger-scale phenomenon possibly linked to a strengthening of
the winter NAO, which might have enhanced the AMOC (Cage and Austin, 2010).
Indeed, several studies report long-lasting warm conditions in Europe
associated with the year 1540 (Casty et al., 2005; Pauling et al., 2006; Wetter
et al., 2014), which given our age model uncertainty (±40 year, see
Table 2) for the 1538–1664 CE time interval may well fall within the warm
period identified for the LIA from our BWT record. A warming around 1540 is
also seen in winter temperature reconstruction for Stockholm ports and
harbours based on historical records of sea ice (Leijonhufvud et al., 2009).
The model-based reconstruction by Orth et al. (2016) suggests that the
European temperatures of 1540 exceeded those of the 2003 summer, which was
likely the warmest for centuries (e.g. Luterbacher et al, 2016). However, this is
difficult to deduce based on the data presented herein, as (i) the
fjord BWT represent winter temperatures and (ii) the record only stretches
until ∼1996.
The climax or the coldest part of the LIA is often linked to the Maunder
minimum in solar activity, which occurred at ∼1645–1715 CE (Mauquoy
et al., 2002). Our record shows a distinct cooling at around 1750 CE with
temperatures ∼1∘C below the 1961–1999 mean, which given a
calibrated 14C age range for this particular date (1675–1813 CE ±25 years; see Table 2), may well represent the Maunder minimum in our
record. At the same time, a 500-year long reconstruction of Stockholm winter
temperatures based on sea ice records from local ports and harbours does not
show the coldest temperatures during the LIA climax, instead demonstrating
that the coldest decade for the last 500 years occurred from 1592 to 1601 CE
with average negative temperature anomalies of ∼-4∘C
(Leijonhufvud et al., 2009).
It is rather intriguing that the coldest bottom water temperatures for
the last 2500 years in the Gullmar Fjord are associated with the onset of the
LIA (1350 CE, ∼2∘C colder than the 1961–1999 mean) rather
than with its climax (Figs. 5–6). This agrees well with the LIA temperature
evolution reported for Loch Sunart (Cage and Austin, 2010) and Chesapeake Bay
(Cronin et al., 2003), which both show 2–4 ∘C cooling of the bottom
waters at the MCA–LIA transition (Fig. 6), attributed to a switch from the
positive winter NAO mode dominating during the medieval times (e.g. Trouet et
al., 2009; Faust et al., 2016) to the negative NAO prevailing during the
major part of the Little Ice Age. Such a switch in the NAO has been linked to
a relaxation of the persistent La-Niña-like conditions in the
equatorial Pacific dominating the MCA (Trouet et al., 2009). The MCA–LIA
transition has been dated to 1250 CE (Cunningham et al., 2013), 1400 CE
(McGregor et al., 2015) and 1450 CE (Cage and Austin, 2010), in contrast to
our study (1350 CE), which may again be a result of 14C dating
uncertainties valid for all of the above-mentioned marine records. At the
same time the Chesapeake Bay study places the MCA–LIA transition between 1300
and 1400 CE (Cronin et al., 2003), which agrees with our data.
Another interesting feature of the LIA climate variability is associated with
consistently low fjord BWT as well as reduced air temperatures during the
1790–1820 CE period as indicated by the Stockholm and central England
instrumental time series (Fig. 8). Despite that this time period is known to coincide with the Dalton
minimum in solar activity (Grove, 1988) it is suggested that volcanic
activity played a much more important role in climate cooling (e.g. Wagner
and Zorita, 2005; McGregor et al., 2015). The role of AMOC strength in
shaping the LIA cold periods is also somewhat controversial based on marine
geological evidence: whilst the AMOC weakening was proposed as a trigger for
the LIA cooling (Bianchi and McCave, 1999), it was argued against (Keigwin
and Boyle, 2000) and was not statistically significant in paleoclimate
modelling (Van der Schrier and Barkmeijer, 2005). It has even been suggested
that the Gulf Stream may have experienced warming during this period (e.g.
Keigwin and Pickart, 1999), which certainly does not explain low BWT
temperatures in our record or low air temperatures over Stockholm and central
England during 1790–1820 CE. An explanation for this phenomenon has been
proposed by Bjerknes (1965), who postulated that “a decrease in the western
European winter surface air temperatures from 1790 to 1820 CE was almost
completely related to the anomalous southward advection of cold polar air”,
a hypothesis later supported by a model study by Van der Schrier and
Barkmeijer (2005).
The Contemporary Warm Period (∼1850 CE–1996)
Most of the proxy records in the North Atlantic indicate a clear warming
trend for the last 100–200 years (Hald et al., 2011, and references therein)
and so do our data, which pick up the warm 1930s and the 1990s (Fig. 8). The 500-year long
reconstruction of Stockholm winter temperatures also demonstrates that the
20th century has experienced four out of the five warmest decades over the
last 500 years: 1905–1914, 1930–1939, 1989–1998 and 1999–2008
(Leijonhufvud et al., 2009). The Gullmar Fjord temperature record shows that
when considering a 3-point running mean temperature variability, the most
recent warming does not stand out in comparison to the RWP and the MCA, as
has also been previously demonstrated by other studies such as a tree
ring-based summer temperature record from central Scandinavia (Linderholm and
Gunnarson, 2005), Scottish loch data (Cage and Austin, 2010), North Atlantic
SST composite data (Cunningham et al., 2013) and a 2000-year temperature
record for continental Europe (PAGES 2K, 2013). At the same time, this “not
outstanding recent warming” seen in our dataset is in contrast to the
Malangen Fjord record (Hald et al., 2011), which infers that the last
100 years have been the warmest in the last two millennia. This may reflect
the so-called polar amplification, as suggested by Hald et al. (2011), since Malangen Fjord is
located much further to the north than Loch Sunart and Gullmar Fjord, which
are both comparably temperate fjord inlets. On the other hand, the stronger recent warming of the Norwegian fjord
record may also be explained by a more direct link to the northward flow of
the Atlantic water as compared to Gullmar Fjord, which is (i) not located
within the core of the North Atlantic Current (Fig. 1) and (ii) reflects
temperature variability during the winter season. At the same time, the
spring SST reconstruction from Chesapeake Bay (Cronin et al., 2003; Fig. 6
herein) shows that the 20th century warming clearly exceeds temperatures
observed during the prior 2500 years. The shallow Chesapeake Bay displays
large seasonal temperature and salinity variability (Cronin et al., 2003) in
contrast to Gullmar Fjord, Malangen Fjord and Loch Sunart, which all have
slightly/ less variable bottom water conditions during the year and similar
“fjordic” circulation with annual or less frequent basin water exchanges.
Furthermore, the SST record from the Chesapeake Bay is the shallowest
temperature reconstruction (12–25 m w.d.) among the temperature records
considered in this study (Loch Sunart: 56 m; Gullmar Fjord: 120 m; and
Malangen Fjord: 218 m w.d.). Shallow water areas are known to generally
warm up faster, especially given the facilitating atmospheric warming of the
late 20th century due to the increase in greenhouse gas emissions (e.g.
Masson-Delmotte, 2013); this may also explain why the recent SST increase in
the Chesapeake Bay record is unprecedented in a 2500-year perspective.
Studying the instrumental hydrographic time series from the Gullmar Fjord
plotted versus reconstructed temperatures (Fig. 7), it is clear that our
record captures the most recent warm period with the bottom water
temperatures that have increased by ∼1.5∘C since the 1960s.
A similar increase has been documented for Loch Sunart (Cage and Austin, 2010)
and Ranafjorden, on the north-west coast of Norway (Klitgaard-Kristensen et al., 2004).
Instrumental meteorological time series for air temperatures since 1960s from
Stockholm and central England also demonstrate a winter temperature
increase of 3–3.5 ∘C, which is higher than the reconstructed range
of Gullmar Fjord bottom water temperatures for this period (Fig. 8). Overall,
the variability in the reconstructed fjord temperatures corresponds well with
both meteorological datasets from 1750 to 1990, with the exception of an individual
wiggle mismatch between 1930 and 1990 (Fig. 8). In general, it appears that both air temperatures records lead the observed
variability while bottom water temperatures are lagging behind for the 1930–1990 period (Fig. 8).
Our record also shows higher BWT prior to the 1920s (Fig. 8), which coincides
with the cold AMO (low SSTs) and low sea surface salinities in the North
Atlantic and subpolar gyre (Reverdin et al. 1994; Reverdin, 2010); however,
in the following period until ∼1960, the reconstructed BWT remains at a
lower level (during the warm AMO, i.e., high North Atlantic SSTs), after
which it peaks again during the “Great Salinity Anomaly” in the late 1970s
and late 1980s (Dickson et al., 1988; Belkin et al., 1998). Nevertheless, it
remains intriguing that on both occasions (prior to the 1920s and during the
1970s/1980s) of reduced salinities and low SSTs in the North Atlantic, our
record is characterized by high fjord deep water temperatures, which is
consistent with increasing air temperatures in instrumental datasets from
Stockholm and central England (Fig. 8). The low surface salinities of the
Great Salinity Anomaly were likely driven by an increased freshwater/sea ice
export from the Arctic via Fram Strait and the Canadian Archipelago (Belkin
et al., 1998). The increased freshwater flux into the subpolar North Atlantic
is, in turn, suggested to have increased the salinity of the North Atlantic
Current, which may have reduced its predicted weakening due to enhanced
freshwater fluxes and would have
helped to restart a stronger AMOC (Hátún et al., 2005; Thornalley et
al., 2009). A stronger North Atlantic Current would subsequently have
resulted in an increased heat transport during winter to the eastern North
Atlantic and in combination with other external forcing factors (e.g. changes
in NAO, volcanism, and solar activity) would have contributed to the warming
observed in the fjord BWT record during the early 20th century. One of those
factors, the positive NAO mode, which has prevailed since the 1970s/1980s
(Hurrell, 1995;
http://www.cpc.ncep.noaa.gov/products/precip/CWlink/pna/season.JFM.nao.gif;
last access: 15 March 2017), extracts heat from the subpolar North Atlantic
through increased westerlies over that region, decreases SSTs, enhances
convection, increases ocean density (Delworth et al., 2016; Delworth and
Zeng, 2016) and results in milder winter conditions over north-western
Europe; thus, it counteracts the effects of the AMOC weakening, which has
been suggested for the 20th century based on modelling data and proxy records
(Caesar et al., 2018; Thornalley et al., 2018). Furthermore, as it is located
within a coastal region, the Gullmar Fjord is more susceptible to wind-forced
temperature changes, which follow the variability of the NAO index and drive
coastal upwelling and downwelling in the fjord (Björk and Nordberg,
2003). According to Jansen et al. (2007), the late 20th century warming, as
demonstrated by many proxy records from the north-eastern Atlantic (see
discussion above), is unlikely to be explained by external forcing factors
and is probably linked to anthropogenic drivers such as greenhouse gas
emissions and aerosols (Booth et al., 2012), which have both significantly
increased since ∼1970s (Masson-Delmotte, 2013).
When studying the Gullmar Fjord bottom water temperature record for the last
2500 years, it is interesting to note that the most recent warming of the
20th century (presented herein until 1996) does not stand
out, and it actually appears to be
comparable to both the Roman Warm Period and the Medieval Climate Anomaly.
However, this observation must be interpreted with caution, as our dataset
does not go beyond year 1996 due to a lack of material (see discussion
below); hence, it does not cover the most recent part of the 20th century
warming, which is widely accepted as having been triggered by growing
anthropogenic emissions.
Environmental conditions explaining the absence or rare occurrence of
Cassidulina laevigata in the record
Since 1990, Cassidulina laevigata has dramatically decreased in
abundance in the Gullmar Fjord deep basin (Fig. 6). A similar pattern, with
short disappearances of C. laevigata, is seen during the Roman and
Medieval warm periods (Fig. 6). This effect may either be due to
increased temperatures and/or, more likely, due to periods of severe hypoxia
as C. laevigata is documented to be sensitive to oxygen
concentrations below 1 mL L-1 (e.g. Gustafsson and Nordberg, 2001;
Nardelli et al., 2014). To a large extent, the oxygen status of fjords and
estuaries on the Swedish west coast, is controlled by climate (e.g. Nordberg
et al., 2000; Filipsson and Nordberg 2004a, b), but the late Holocene changes
in land use and organic enrichment in the fjord are also suggested to have played a
role (Filipsson and Nordberg, 2010). Thus, the short extinctions of
C. laevigata during warmer periods in the past may be
equivalent to the present-day pattern of severe hypoxia following the
positive NAO periods with mild and humid winters, limited basin water
exchange and high organic matter flux increasing oxygen demand (Nordberg et
al., 2000, 2001; Filipsson and Nordberg 2004a). Indeed, when comparing our
record to the reconstructed NAO index from the Trondheim Fjord in western Norway
(Faust et al., 2016) it appears that sediment core intervals with absent
C. laevigata (at ∼75, 450, 1000 CE and post-1990) correlate
rather well with the positive NAO index (Fig. 6).
Conclusions
To conclude, from the available paleotemperature equations, the equation by
McCorkle et al. (1997) produced the most realistic reconstructed deep water
temperature range of 2.7–7.8 ∘C, which falls within the annual
variability instrumentally recorded in the deep fjord basin since 1890. This
suggests that the Gullmar Fjord δ18O record mainly reflects the
variability of the winter bottom water temperatures with a minor salinity
influence. The relationship between the evolution of the fjord's bottom water
temperatures over the last two millennia and other late Holocene climate
records reveals synchronous North Atlantic-wide centennial and multidecadal
climate variability despite age model uncertainties, different proxy type,
time resolution, annual versus seasonal signal and different hydrographic
characteristics.
The record shows a substantial and long-term warming during the Roman Warm
Period (∼350 BCE–450 CE), followed by variable bottom water
temperatures during the Dark Ages (∼450–850 CE). The Viking
Age/Medieval Climate Anomaly (∼850–1350 CE) is also indicated by
positive bottom water temperature anomalies, while the Little Ice Age (∼1350–1850 CE) is characterized by a long-term cooling with distinct
multidecadal variability. The record also picks up the contemporary warming
of the 1930s and the 1990s. When studying the Gullmar Fjord bottom water
temperature record for the last 2500 years, it is interesting to note that
the warming of the 20th century (presented herein until 1996) is comparable
to both the Roman Warm Period and the Medieval Climate Anomaly.
Data availability
The data presented in this paper are available at PANGEA
(https://doi.org/10.1594/PANGAEA.892500).
Information about the Supplement
Figure S1 contains a scatter plot of stable carbon isotope (δ13C) data from the composite G113-2Aa – 9004 record (Filipsson and
Nordberg, 2010) plotted against the oxygen isotope data presented herein.
Note absence of a correlation between the two, ruling out the possibility that
the changes in δ18O are due to changes in water masses.
The supplement related to this article is available online at: https://doi.org/10.5194/cp-14-1097-2018-supplement.
Author contributions
KN conceived the research, obtained funding and organized and
performed sediment core sampling in 1990 and 1999. HLF participated in the
1999 cruise, picked most of the foraminiferal samples and prepared them for
stable oxygen isotopes and funded isotope analysis. IPA participated in an
additional sampling campaign in 2009, undertook sediment core sampling and picked and
prepared the samples for stable isotope analysis. All authors (IPA, KN and
HLF) equally contributed to the data analysis and interpretation. IPA wrote the
manuscript with the help of both co-authors.
Competing interests
The authors declare that they have no conflict of
interest.
Acknowledgements
The authors sincerely thank everyone who helped perform this study. The crews
of R/V Svanic, R/V Arne Tiselius and R/V Skagerak
assisted with sampling. This study was financed by the following
institutions: the Swedish Research Council (KN) – grants no. 621-2004-5320
and 621-2007-4369, Swedish Research Council (HLF) – grant no. 621-2005-4265;
the Lamm Foundation (KN); the Marine Research Centre, GMF (KN); and EUROPROX
– European Graduate College – Proxies in Earth History (HLF). Monika Segl
(University of Bremen) measured the stable O and C isotopes. The PALEOSTUDIES
program (University of Bremen) covered the costs for isotope analyses, while
the Department of Earth Sciences (University of Gothenburg) provided a
postdoctoral fellowship to IPA. The manuscript greatly benefited from the
insightful comments and suggestions from Antoon Kuijpers, an anonymous
reviewer and journal editor Alessio Rovere.
The hydrographic data used in the study were obtained from the SMHI
oceanographic observation database (SHARK). The SHARK data collection is
organized by the environmental monitoring program and is funded by the
Swedish Agency for Marine and Water Management (SwAM). Edited by: Alessio Rovere Reviewed by:
Marit-Solveig Seidenkrantz and Antoon Kuijpers
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