Introduction
The Mediterranean is considered one of the most vulnerable regions
with regard to the current global warming (Giorgi, 2006). This high
sensitivity to climate variability has been evidenced in several studies on
past natural changes (Rohling et al., 1998; Cacho et al., 1999a; Moreno et
al., 2002; Martrat et al., 2004; Reguera, 2004; Frigola et al., 2007;
Combourieu Nebout et al., 2009). Palaeo-studies focussed mostly on the rapid
climate variability in the last glacial period have shown solid evidence of
a close connection between changes in North Atlantic oceanography and
climate over the western Mediterranean region (Cacho et al., 1999b, 2000,
2001; Moreno et al., 2005; Sierro et al., 2005; Frigola et al., 2008;
Fletcher and Sanchez-Goñi, 2008). Nevertheless, climate variability
during the Holocene, and particularly during the last millennium, is not so
well described in this region, although its understanding is crucial for
placing the nature of the 20th century trends in the recent climate history
(Huang, 2004).
Some previous studies have already proposed that the Holocene centennial
climate variability in the western Mediterranean Sea could be linked to the
North Atlantic Oscillation (NAO) variability (Jalut et al., 1997, 2000; Combourieu Nebout et al., 2002;
Goy et al., 2003; Roberts et al., 2012; Fletcher et al., 2012). In
particular, nine Holocene episodes of enhanced deep water convection in the
Gulf of Lion (GoL) and surface cooling conditions have been described in the
region (Frigola et al., 2007). These events have also been correlated to
intensified upwelling conditions in the Alboran Sea and tentatively
described as two-phase scenarios driven by distinctive NAO states
(Ausín et al., 2015).
A growing number of studies have revealed considerable climate fluctuations during
the last 2 kyr (Abrantes et al., 2005; González-Álvarez et al.,
2005; Holzhauser et al., 2005; Kaufman et al., 2009; Lebreiro et al., 2006;
Martín-Puertas et al., 2008; Pena et al., 2010; Kobashi et al., 2011;
Nieto-Moreno et al., 2011, 2013; Moreno et al., 2012; PAGES 2K Consortium,
2013; Esper et al., 2014; McGregor et al., 2015). However, there is no
agreement on the exact time span of the different climatic periods defined,
such as the Medieval Climatic Anomaly (MCA), a term coined
originally by Stine (1994).
The existing Mediterranean climatic records for the last 1 or 2 kyr are
mostly based on terrestrial archives such as tree rings (Touchan et al.,
2005, 2007; Griggs et al., 2007; Esper et al., 2007; Büntgen et al., 2011; Morellón et al., 2012), speleothem records (Frisia
et al., 2003; Mangini et al., 2005; Fleitmann et al., 2009;
Martín-Chivelet et al., 2011; Wassenburg et al., 2013), or lake
reconstructions (Pla and Catalan, 2005; Martín-Puertas et al., 2008;
Corella et al., 2011; Morellón et al., 2012). All of these archives can
be good sensors of temperature and humidity changes, but it is often
difficult to disentangle the effect of both variables in the proxy records.
Recent efforts focussed on integrating these 2 kyr records into regional
climatic signals reveal complex regional responses and evidence the scarcity
of marine records to form a more complete picture (PAGES, 2009;
Lionello, 2012).
Marine records are often limited by the lack of
adequate time resolution and/or robust chronologies for detailed comparison
with terrestrial records. However, marine records provide a wider
range of temperature-sensitive proxies. There are few marine palaeoclimate
records available for the last 2 kyr in the Mediterranean Sea (Schilman et
al., 2001; Versteegh et al., 2007; Piva et al., 2008; Taricco et al., 2009,
2015; Incarbona et al., 2010; Fanget et al., 2013; Grauel et al., 2013;
Lirer et al., 2013, 2014; Di Bella et al., 2014; Goudeau et al., 2015) and
they are even more scarce in the western basin. Unfortunately, the existing
pool of marine proxy data in the Mediterranean for the last two millennia is
too sparse to recognize common patterns of climate variability (Taricco et
al., 2009; Nieto-Moreno et al., 2011; Moreno et al., 2012, and references
therein).
The aim of the present study is to characterize changes in surface water
properties from the Minorca margin in the Catalan–Balearic Sea
(central-western Mediterranean) in order to contribute to a better understanding of
the climate variations in this region during the last 2.7 kyr. Sea surface
temperature (SST) has been reconstructed by means of two independent
proxies, Mg / Ca analyses on the planktonic foraminifera Globigerina bulloides and alkenone-derived
SST (Villanueva et al., 1997; Lea et al., 1999; Barker et al., 2005; Conte
et al., 2006). The application of G. bulloides Mg / Ca as a palaeothermometer in the
western Mediterranean Sea is tested through the analysis of a series of core-top samples from different locations of the western Mediterranean Sea and
the calibration reviewed consistently. Mg / Ca thermometry is applied in
conjunction with δ18O in order to evaluate changes in the
evaporation–precipitation (E–P) balance of the basin, which are ultimately
linked to salinity (Lea et al., 1999; Pierre, 1999; Barker et al., 2005).
Location of the studied area. (a) Central-western Mediterranean
Sea: cores MIN and MR3 (red dots). NC: Northern Current (surface). WMDW:
Western Mediterranean Deep Water. (b) Cores used in age model development
from the Tyrrhenian Sea (green triangles; Lirer et al., 2013) and cores
used in Mg / Ca–SST calibration from the western Mediterranean Basin (blue
squares).
One of the intrinsic limitations of studying the climate evolution of the
last 2 kyr is that the magnitude of climatic oscillations is often below the
sensitivity of the selected proxies. In order to overcome this limitation we
have produced “stack” proxy records from multicores in the same region. The
stack record captures the first-order climatic variability from the proxy
records and removes the noise, therefore allowing for a more robust
identification of regional climatic variability.
The studied time periods have been defined as follows (years expressed as
BCE, before common era, and CE, common era): the Talaiotic Period (TP, ending
in 123 BCE), Roman Period (RP, from 123 BCE to 470 CE), “Dark Middle
Ages” (DMA; from 470 until 900 CE), Medieval Climate Anomaly (MCA, from 900 to
1275 CE), and Little Ice Age (LIA, from 1275 to 1850 CE), with the Industrial Era (IE)
as the most recent period. The limits of these periods are not uniform
across the Mediterranean (Lionello, 2012), and here the selected ages have
been chosen according to historical events in Minorca and to the
classic climatic ones defined in the literature (i.e. Nieto-Moreno et al.,
2011, 2013; Moreno et al., 2012; Lirer et al., 2013, 2014).
Climatic and oceanographic settings
The Mediterranean Sea is a semi-enclosed basin located in a transitional
zone between different climate regimes, from the temperate zone in the
north to the subtropical zone in the south. Consequently, the Mediterranean
climate is characterized by mild wet winters and warm to hot, dry summers
(Lionello et al., 2006). Interannual climate variability is very much
controlled by the dipole-like pressure gradient between the Azores (high)
and Iceland (low) system, known as the North Atlantic Oscillation (NAO; Hurrell, 1995; Lionello and Sanna, 2005; Mariotti, 2011; Ausín et al.,
2015). However, the northern part of the Mediterranean region is also linked to
other mid-latitude teleconnection patterns (Lionello, 2012).
The Mediterranean Sea is a concentration basin (Béthoux, 1980; Lacombe
et al., 1981) and the excess of evaporation with respect to freshwater input
is balanced by water exchange at the Strait of Gibraltar (i.e. Pinardi and
Masetti, 2000; Malanotte-Rizzoli et al., 2014). The basin-wide circulation
pattern is predominantly cyclonic (Millot, 1999). Three convection cells
promote the Mediterranean deep and intermediate circulation: a basin-wide
open cell and two separated closed cells, one for the western part of the basin and one
for the eastern part. The first one connects the two basins of the
Mediterranean Sea though the Strait of Sicily, where water masses interchange
occurs at intermediate depths. This cell is associated with the inflow of
Atlantic Water (AW) at the Strait of Gibraltar and the outflow of the
Levantine Intermediate Water (LIW) that flows below the first (Lionello et
al., 2006).
In the north-western Mediterranean Sea, the Northern Current (NC) represents
the main feature of the surface circulation transporting waters alongshore
from the Ligurian Sea to the Alboran Sea (Fig. 1a). North-east of the
Balearic Promontory a surface oceanographic front separates Mediterranean
waters transported by the NC from the Atlantic waters that recently entered
the Mediterranean (Millot, 1999; Pinot et al., 2002; André et al.,
2005).
Deep convection occurs offshore of the GoL due to the action of persistent cold
and dry winter winds such as the tramontana and the mistral. These winds
cause strong evaporation and cooling of surface water, thus increasing their
density, sinking to greater depths and leading to Western Mediterranean Deep
Water formation (WMDW; MEDOC, 1970; Lacombe et al., 1985; Millot, 1999).
Dense shelf water cascading (DSWC) in the GoL also contributes to the sink
of large volumes of water and sediments into the deep basin (Canals et al.,
2006).
The north-western Mediterranean primary production is subject to an intense
bloom in late winter–spring, when the surface layer stabilizes, and sometimes
to a less intense bloom in autumn, when the strong summer thermocline is
progressively eroded (Estrada et al., 1985; Bosc et al., 2004; D'Ortenzio
and Ribera, 2009; Siokou-Frangou et al., 2010). SST in the region evolves
accordingly with the seasonal bloom, with minima SST in February, which
subsequently increases until maximum SST values during August. Afterwards, a
SST drop can be observed in October, although with some interannual
variability (Pastor, 2012).
Material and methods
Sediment core description
The studied sediment cores were recovered from a sediment drift built by the
action of the southward branch of the WMDW north of Minorca (Fig. 1).
Previous studies carried out at this site have already described high
sedimentation rates (> 20 cm kyr-1; Frigola et al., 2007,
2008; Moreno et al., 2012), suggesting that this location was suitable
for a detailed study of the last millennia. The cores were recovered with a
multicore system in two different stations located at about 50 km north of
Minorca. Cores MINMC06-1 and MINMC06-2 (henceforth MIN1 and MIN2; 40∘29′ N, 04∘01′ E; 2391 m water depth; 31 and 32.5 cm
core length, respectively) were retrieved in 2006 during the HERMES 3 cruise
onboard the R/V Thethys II. The recovery of cores HER-MC-MR3.1,
HER-MC-MR3.2, and HER-MC-MR3.3 (henceforth MR3.1, MR3.2, and MR3.3; 40∘29′ N,
3∘37′ E; 2117 m water depth; 27, 18, and 27 cm core length,
respectively) took place in 2009 during the HERMESIONE expedition onboard
the R/V Hespérides. The distance between MIN and MR3 cores is
∼ 30 km and both stations are located at an intermediate
position within the sediment drift, which extends along a water depth range
from 2000 to 2700 m (Frigola, 2012; Velasco et al., 1996; Mauffret,
1979). The MIN cores are from sites that are about 300 m deeper than
the MR3 ones.
Core tops included in the calibration's adjustment. δ18Oc and Mg / Ca have been
obtained from G. bulloides (Mg / Ca procedure has
been performed without reductive step).
Core
Location
Latitude
Longitude
Mg / Ca
δ18Oc
(mmol mol-1)
(‰ VPDB)
TR4-157
Balearic Abyssal Plain
40∘30.00′ N
4∘55.76′ E
3.36
0.53
ALB1
Alboran Sea (W. Med.)
36∘14.31′ N
4∘15.52′ W
3.20
0.80
ALBT1
Alboran Sea (W. Med.)
36∘22.05′ N
4∘18.14′ W
3.44
0.65
ALBT2
Alboran Sea (E. Med.)
36∘06.09′ N
3∘02.41′ W
3.63
0.57
ALBT4
Alboran Sea (E. Med.)
36∘39.63′ N
1∘32.35′ W
3.72
0.93
ALBT5
Alboran Sea (E. Med.)
36∘13.60′ N
1∘35.97′ W
3.38
0.64
MIN cores were homogeneously sampled at 0.5 cm resolution in the laboratory.
In the MR3 cores a different strategy was followed. MR3.1 and MR3.2 were
initially subsampled with a PVC tube and split into two halves for X-ray fluorescence (XRF)
analyses in the laboratory. Both halves of core MR3.1 (MR3.1A and MR3.1B)
were used for the present work as replicates of the same core, and records
for each half are shown separately. All MR3 cores were sampled at 0.5 cm
resolution in the upper 15 cm and at 1 cm in the deeper sections, with the
exception of MR3.1B that was sampled at 0.25 cm resolution. The MR3
cores were composed of brown–orange nannofossil and foraminifera silty clay,
which was lightly bioturbated and contained layers enriched in pteropods and
fragments of gastropods as well as some dark layers.
Additionally, core-top samples from seven multicores collected at different
locations in the western Mediterranean have also been used for the
correction of the Mg / Ca–SST calibration from G. bulloides (Table 1; Fig. 1).
Radiocarbon analyses
Twelve 14C AMS dates were measured in cores MIN1, MIN2, and MR3.3
(Supplement Table S1) using 4–22 mg samples of the planktonic foraminifer
Globorotalia inflata handpicked from the > 355 µm fraction. The ages were
calibrated with the standard marine correction of 408 years and the regional
average marine reservoir correction (ΔR) for the central-western
Mediterranean Sea using Calib 7.0 software (Stuiver and Reimer, 1993) and
the MARINE13 calibration curve (Reimer et al., 2013).
Radionuclides 210Pb and 137Cs
The concentrations of the naturally occurring radionuclide 210Pb
(Supplement Fig. S1) were determined in cores MIN1, MIN2, MR3.1A, and
MR3.2 by alpha spectroscopy (Sanchez-Cabeza et al., 1998). The concentrations
of the anthropogenic radionuclide 137Cs in core MIN1 (Fig. S1) were measured by gamma spectrometry using a high-purity intrinsic
germanium detector. The 226Ra concentrations were determined from the
gamma emissions of 214Pb that were also used to calculate the excess
210Pb concentrations. The sediment accumulation rates for the last
century (Sect. S1.1 in the Supplement) were calculated using the CIC (constant
initial concentration) and the CF : CS (constant flux : constant
sedimentation) models (Appleby and Oldfield, 1992; Krishnaswami et al.,
1971), constrained by the 137Cs concentration profile in core MIN1
(Masqué et al., 2003).
Bulk geochemical analyses
The elemental composition of cores MR3.1B and MR3.2 was obtained with an Avaatech XRF
core-scanner system (CORELAB, University of Barcelona), which is
equipped with an optical variable system that allows determining the length
(10–0.1 mm) and the extent (15–2 mm) of the bundle of X-rays in an
independent way. This method allows obtaining qualitative information of the
elementary composition of the core materials. The core surfaces were
scraped, cleaned, and covered with a 4 µm thin SPEXCertiPrep Ultralene foil to
prevent contamination and minimize desiccation (Richter and van der Gaast,
2006). Sampling was performed every 1 cm and scanning took place at the
split core surface directly. Among the several elements measured in this
study, the Mn profile was used for the construction of the age models (see
Supplement for age model development).
Planktonic foraminiferal analyses
Planktonic foraminifera specimens of Globigerina bulloides were picked together from a size range
of 250–355 µm, crushed, and cleaned separately for Mg / Ca and δ18O measurements. In core MR3.1B, picking was often performed in the
< 355 µm fraction due to the small amount of material (sampling
every 0.25 cm). Additionally, quantitative analysis of planktonic
foraminiferal assemblages was carried out in core MR3.3 and on the upper
part of core MR3.1A by using the fraction size above 125 µm
(Fig. S2). The 42 studied samples showed abundant and
well-preserved planktonic foraminifera.
The samples for trace elements analyses consisted of ∼ 45 specimens of G. bulloides that were
crushed under glass slides to open the chambers.
Foraminifera cleaning consisted of clay removal and oxidative and weak acid
leaching steps (Pena et al., 2005). Samples from core MR3.1A were also
cleaned including the “reductive step”. Elemental ratios were measured on
an inductively coupled plasma mass spectrometer (ICP-MS, Perkin Elmer ELAN
6000) in the Scientific and Technological Centers of the University of
Barcelona (CCiT-UB). A standard solution with known elemental ratios was
used for sample standard bracketing (SSB) as a correction for instrumental
drift. The average reproducibility of Mg / Ca ratios, taking into account the
known standard solutions concentrations, was 97 and 89 % for MIN1 and MIN2
cores, and 99 and 97 % for cores MR3.1A, MR3.1B, and MR3.3, respectively.
Procedural blanks were routinely measured to detect any potential
contamination problem during cleaning and dissolution. The Mn / Ca and Al / Ca
ratios were also always measured to identify potential contaminations due to
the presence of manganese oxides and/or aluminosilicates (Barker et al., 2003; Lea
et al., 2005; Pena et al., 2005, 2008).
To avoid the overestimation of Mg / Ca–SST by diagenetic contamination, Mn / Ca
values > 0.5 mmol mol-1 were discarded from core MR3.1B and
only those higher than 1 mmol mol-1 were removed from MIN1 and MR3.3. Samples
suspected to have detrital contamination with elevated Al / Ca ratios were also
removed. No significant correlation exists between Mg / Ca and Mn / Ca or Al / Ca
ratios after data filtering (r < 0.29, p value = 0.06).
The Mg / Ca ratios were transferred into SST values using the calibration
proposed in the present study (Sect. 5.1). In the case of the MR3.1A record,
which was cleaned using the reductive procedure, and as was expected (Barker et al.,
2003; Pena et al., 2005; Yu et al., 2007), the Mg / Ca ratios were about
23 % lower than those measured in core MR3.1B without the reductive step.
The obtained percentage of Mg / Ca lowering is comparable to or higher than
those percentages previously estimated for different planktonic foraminifera, although data
from G. bulloides have not been previously reported (Barker et al., 2003). Mg / Ca–SST in core
MR3.1A was calculated after the Mg / Ca correction of this 23 % offset by
application of the calibration used with the other records.
Stable isotope measurements were performed by means of sonication on 10 specimens of G. bulloides after
methanol cleaning to remove fine-grained particles. The
analyses were performed in a Finnigan MAT 252 mass spectrometer fitted with
a Kiel-IV carbonate microsampler in the CCiT-UB. The analytical precision of
laboratory standards for δ18O was better than 0.08 ‰. Calibration to Vienna Pee Dee Belemnite (VPDB) was
carried out by means of NBS-19 standards (Coplen, 1996).
Seawater δ18O (δ18Osw) was obtained after
removing the temperature effect on the G. bulloides δ18O signal using the
Mg / Ca–SST records of the Shackleton palaeotemperature equation (Shackleton,
1974). The results are expressed in the SMOW (Standard Mean Ocean Water) water standard (δ18Osw) after the correction of Craig (1965). The use of specific
temperature equations for G. bulloides was also considered (Bemis et al., 1998; Mulitza
et al., 2003), but the core-top estimates provided
δ18Osw
values of 2.1–1.5 ‰ SMOW, which were significantly
higher than those measured in water samples from the central-western Mediterranean Sea (∼ 1.2 ‰ SMOW) (Pierre, 1999).
After application of the empirical Shackleton (1974) palaeotemperature
equation, the core-top δ18Osw estimates averaged
1.1 ‰ SMOW and were closer to the actual seawater
measurements. This, it was decided that this equation provided more
realistic oceanographical conditions in this location.
Alkenones
Measurements of the relative proportion of unsaturated C37 alkenones,
namely the U37k′ index, were carried out in order to obtain SST
records for the studied cores. Detailed information about the methodology and
equipment used can be found in Villanueva et al. (1997). The precision of
this palaeothermometry tool has been determined to be about ±0.5 ∘C (Eglinton et al., 2001). Furthermore, taking into account
duplicate alkenone analysis carried out on core MR3.3, the precision
achieved results better than ±0.8 ∘C. The reconstruction of
SST records was based on the global calibration of Conte et al. (2006),
which considers an estimation standard error of 1.1 ∘C in surface
sediments.
Sea surface temperatures and δ18O data
Mg / Ca–SST calibration
The Mg / Ca ratio measured in G. bulloides is a widely used proxy to reconstruct SST
(Barker et al., 2005), although the calibrations available can provide very
different results
(Lea et al., 1999; Mashiotta et al., 1999; Elderfield and Ganssen, 2000;
Anand et al., 2003; McConnell and Thunell, 2005; Cléroux et al., 2008;
Thornalley et al., 2009; Patton et al., 2011). Apparently, the regional
Mg / Ca–temperature response varies due to parameters that have not yet been
identified (Patton et al., 2011). A further difficulty arises from the
questioned Mg / Ca thermal signal in high-salinity regions such as the
Mediterranean Sea, where anomalously high Mg / Ca values have been observed
(Ferguson et al., 2008). This apparent high salinity sensitivity in
foraminiferal Mg / Ca ratios is under discussion and has not been supported
by recent culture experiments (Hönisch et al., 2013), which, in addition,
could be attributed to diagenetic overprints (Hoogakker et al., 2009; van
Raden et al., 2011). In order to test the value of the Mg / Ca ratios in G. bulloides from
the western Mediterranean Sea and also review its significance in terms of
seasonality and depth habitat, a set of core-top samples from different
locations of the western Mediterranean Sea have been analysed. Core-top
samples were recovered using a multicorer system, and they can be considered
representative of present conditions (Masqué et al., 2003;
Cacho et al., 2006). The studied cores are located in the 35–45∘ N latitude range (Table 1 and Fig. 1) and mostly represent two different
trophic regimes, defined by the classical spring bloom (the most
north-western basin) and an intermittent bloom (D'Ortenzio and Ribera,
2009).
(a) Exponential function and correlation between δ18Oc temperatures and Mg / Ca. Dashed lines show the 1σ
confidence limits of the curve fit. The standard error of our temperature
calibration taking into account each δ18Oc temperature
from core tops (Table 1) is ±0.6 ∘C. Error of temperature
estimates based on our G. bulloides calibration for the western Mediterranean is
±1.4 ∘C. These uncertainties are higher but still in the range of
±0.6 ∘C obtained for the Atlantic Ocean in Elderfield and
Ganssen (2000) and also 1.1 ∘C in G. bulloides culture data (Lea
et al., 1999). (b) April (red) and May (black) temperature profiles of the
first 200 m measured during years 1945–2000 in stations corresponding to the
studied core tops (MEDAR GROUP, 2002). The δ18Oc average
temperature of all cores is shown (grey, vertical band).
The resulting Mg / Ca ratios have been compared with the isotopically derived
calcification temperatures based on the δ18O measurements
performed also in G. bulloides from the same samples. This comparison was performed
after use of the Shackleton (1974) palaeotemperature equation and the δ18Owater data published by Pierre (1999), always considering the
values of the closer stations and the top 100 m. The resulting Mg / Ca–SST
data have been plotted together with those of G. bulloides from North Atlantic core tops
previously published by Elderfield and Ganssen (2000). The resulting high
correlation (r2= 0.92; Fig. 2a) strongly supports that the Mg / Ca
ratios of the central-western Mediterranean Sea are dominated by a thermal
signal. Thus, the new data set from the Mediterranean core tops improves
temperature sensitivity range over the warm end of the calibration. The
resulting exponential function indicates ∼ 9.4 % Mg / Ca per
∘C sensitivity in the Mg uptake with respect to temperature,
which is in agreement with the range described in the literature (i.e.,
Elderfield and Ganssen, 2000; Barker et al., 2005; Patton et al., 2011). The
new equation for the Mg / Ca–SST calibration including data from the western
Mediterranean Sea and the Atlantic Ocean is as follows:
Mg/Ca=0.7045(±0.0710)e0.0939(±0.0066)T.
The Mg / Ca–SST signal of G. bulloides has been compared with a compilation of water
temperature profiles of the first 200 m measured between years 1945 and 2000 in
stations close to the studied core tops (MEDAR GROUP, 2002). Although
significant regional and interannual variations have been observed, the
obtained calcification temperatures of our core-top samples show the best
agreement with temperature values of the upper 40 m during the spring months
(April–May; Fig. 2b). This water depth is consistent with preferential
depth ranges for G. bulloides found by plankton tows in the Mediterranean (Pujol and
Vergnaud-Grazzini, 1995) and with results from multiannual sediment trap
monitoring in the Alboran Sea and the GoL, where maximum G. bulloides percentages were
observed just before the beginning of thermal stratifications (see
Bárcena et al., 2004; Bosc et al., 2004; Rigual-Hernández et al.,
2012). Although the information available about depth and seasonality
distribution of G. bulloides is relatively fragmented, this species is generally
found in intermediate or even shallow waters (i.e. Bé and Hutson, 197; Ganssen
and Kroon, 2000; Schiebel et al., 2002; Rogerson et al., 2004; Thornalley et
al., 2009). However, G. bulloides has also been observed at deeper depths in some
western Mediterranean Sea sub-basins (Pujol and Vergnaud-Grazzini, 1995).
Extended data with enhanced spatial and seasonal coverage are required in
order to better characterize production, seasonality, and geographic and
distribution patterns of live foraminifera such as G. bulloides. Nevertheless, the
obtained core-top data set offers solid evidence on the seasonal character
of the recorded temperature signal in the Mg / Ca ratio.
SST obtained from Mg / Ca for cores: (a) MR3.1B, (b) MR3.1A,
(c) MR3.3, (d) MIN2, and (e) MIN1. The grey shaded areas integrate uncertainties of
average values and represent 1σ of the absolute values. This
uncertainty includes analytical precision and reproducibility and the
uncertainties derived from the G. bulloides core-top calibrations for the
central-western Mediterranean Sea developed in this paper. (f) All
individual SST anomalies on their respective time step (MR3.1B: orange;
MR3.1A: purple; MR3.3: green; MIN2: blue; MIN1: black dots). (g) 20 yr cm-1 stacked temperature anomaly (red plot) with its 2σ
uncertainty (grey band). The 80 yr cm-1 (grey plot) and the 100 yr cm-1 (black plot) stacks are also shown. The triangles represent
14C dates (black) and biostratigraphical dates based on planktonic
foraminifera (blue) and are shown below the corresponding core, including
their associated 2σ errors.
A regional stack for Mg / Ca–SST records
The Mg / Ca–SST profiles obtained from our sediment records are plotted with
the resulting common age model in Fig. 3. The average SST values for the
last 2700 years ranged from 16.0 ± 0.9 to 17.8 ± 0.8 ∘C (uncertainties of average values represent 1σ; uncertainties of
absolute values include analytical precision and reproducibility and also
those derived from the Mg / Ca–SST calibration). SST records show the warmest
sustained period during the RP, approximately between 170 BCE and 300 CE, except in core MIN2, since this record ends at the RP–DMA transition. In addition, all the records show a generally
consistent cooling trend after the RP with several centennial-scale
oscillations. The maximum SST value is observed in core MR3.3 (19.6 ± 1.8 ∘C) during the MCA (Fig. 3c) and
the minimum is recorded in core MIN1 (14.4 ± 1.4 ∘C) during
the LIA (Fig. 3e). Centennial-scale variability is
predominant throughout the records. Particularly, during MCA some warm
episodes reached slightly higher SST than the averaged SST maximum (i.e.
19.6 ± 1.8 ∘C at ∼ 1021 CE). These events
were far shorter in duration compared to the RP (Fig. 3). The highest frequency
of intense cold events occurred during the LIA and, in particular, the last
millennium recorded the minimum average Mg / Ca–SST (15.2 ± 0.8 ∘C). Four of the five records show a pronounced SST drop after
1275 CE, coinciding with the onset of the LIA. Based on the different
Mg / Ca–SST patterns, the LIA period has been divided into two subperiods, an
early warmer interval (LIAa) and a later colder interval (LIAb) by reference
to the 1540 CE boundary.
Oxygen isotope measured on carbonate shells of G. bulloides (δ18Oc ‰ VPDB, in black) and their derived
δ18Osw (purple) for cores: (a) MR3.1B, (b) MR3.1A,
(c) MR3.3, (d) MIN2, and (e) MIN1. (f) Individual δ18Oc
(‰ VPDB) anomalies on their respective time step. (g) δ18Oc and δ18Osw anomaly stacked records
(‰ VPDB and ‰ SMOW, respectively).
One of the main difficulties with SST reconstructions in the last millennia is
the internal noise of the records due to sampling and proxy limitations,
which is of the same amplitude as the targeted climatic signal variability.
In this sense, we have constructed a Mg / Ca–SST anomaly stack with the aim to
detect the most robust climatic structures along the different records and
reduce the individual noise. First, each SST record was converted into a SST
anomaly record in relation to its average temperature (Fig. 3f). Secondly,
in order to obtain a common sampling interval, all records were interpolated.
Interpolation at three different resolutions did not result in significant
differences (Fig. 3g). Subsequently, we selected the stack that provided the
best resolution offered by our age models (20 yr cm-1) since it very well
preserves the high-frequency variability of the individual records
(Fig. 3g). The obtained SST anomaly stack allows for a better identification
of the most significant features at centennial timescales. Abrupt cooling
events are mainly recorded during the LIA (-0.5 to
-0.7 ∘C 100 yr-1), while abrupt warmings (0.9 to 0.6 ∘C 100 yr-1) are
detected during the MCA. Events of similar magnitude have been also
documented during the LIA–IE transition. When considering the entire SST
anomaly record, a long-term cooling trend of about -1 to -2 ∘C is
observed. However, focussing on the last 1800 years, since the RP maxima, the
observed cooling trend was far more intense, at about -3.1 to -3.5 ∘C (-0.3 to -0.8 ∘C kyr-1). This is consistent
within the recent 2 kyr global reconstruction published by McGregor et al. (2015; estimation of the SST cooling trend, using the average anomaly
method 1 for the period 1–2000 CE: -0.3 to
-0.4 ∘C kyr-1).
Oxygen isotope records
The oxygen isotopes measured on carbonate shells of G. bulloides (δ18Oc) and their derived δ18Osw after removing
the temperature effect with the Mg / Ca–SST signals (see Sect. 3.5) are shown
in Fig. 4. δ18Oc and their derived δ18Osw profiles have been respectively stacked following the same
procedure for the Mg / Ca–SST stack (Sect. 5.2). In general terms, all the
records present a highly stable pattern during the whole period with a weak
depleting trend, which is almost undetectable in some cases (i.e. core
MIN1).
The average δ18Oc values range from 1.2 to
1.4 ‰ VPDB and, in general, the MR3 cores show slightly
higher values (∼ 1.4 ‰ VPDB) than the MIN
cores (∼ 1.2 ‰ VPDB). The lowest δ18Oc values (1.0 to 1.2 ‰ VPDB) mostly occur
during the RP, although some short, low excursions can also be observed
during the end of the MCA and/or the LIA. The highest values (1.4 to
1.8 ‰ VPDB) are mainly associated with short events
during the LIA, the MCA, and over the TP–RP transition. A significant
increase in δ18Oc values is observed at the LIA–IE
transition, although a sudden drop is recorded at the end of the stack
record (after 1867 CE), which could result from a differential influence
of the records (i.e. MIN1) and/or an extreme artefact (Fig. 4g).
After removing the temperature effect from the δ18Oc
record, the remaining δ18Osw record mainly reflects
changes in E–P balance, thus resulting in an indirect proxy of sea surface
salinity. The average δ18Osw values obtained for the
period studied ranged from 1.3 to 1.8 ‰ SMOW. The
highest δ18Osw values (from 2.4 to
1.9 ‰ SMOW) are recorded during the RP, when the longest
warm period is also observed, and some values are notable during the MCA too.
Enhancements of the E–P balance (δ18Osw higher values)
coincide with higher SST. The lowest δ18Osw values (from
0.8 to 1.5 ‰ SMOW) are recorded particularly during the
onset and the end of the LIA and also during the MCA. A drop in the E-P
balance has been obtained from approximately the end of LIA to the most
recent years. The most significant changes in our δ18Osw
stack record correspond to increases in the most recent times and around
1200 CE (MCA) and to the decrease observed at the end of the LIA (Fig. 4).
Alkenone temperature records from Minorca (this study) for cores
(a) MR3.3, (b) MIN2, and (c) MIN1. Triangles represent 14C dates
(black) and biostratigraphical dates based on planktonic foraminifera
(blue) and are shown below the corresponding core with their associated
2σ errors. (d) Individual alkenone-derived SST anomalies in their
respective time step (MR3.3: green; MIN2: blue; MIN1: black dots).
(e) 20 yr cm-1 stacked temperature anomaly (orange plot). The 80 yr cm-1
(grey plot) and the 100 yr cm-1 (black plot) stacks are also shown.
Alkenone–SST records
The two alkenone (U37k′)-derived SSTs of MIN cores have
already been published in Moreno et al. (2012), while the records from MR3 cores
are new (Fig. 5). The four alkenone–SST records show a similar general
cooling trend during the studied period and they have also been integrated
in a SST anomaly stack (Fig. 5e). The general cooling trend involves about
-1.4 ∘C when the entire studied period is considered and about
-1.7 ∘C since the SST maximum recorded during the RP. The mean SST
uncertainties in this section have been estimated as ±1.1 ∘C, taking into account the estimated standard error (see Sect. 3.6).
Previous studies have interpreted the alkenone–SST signal in the western
Mediterranean Sea as an annual average (Ternois et al., 1996; Cacho et al.,
1999a, b; Martrat et al., 2004). The average alkenone–SST values for the
studied period (last 2700 years) ranged from 17.0 to 17.4 ∘C.
The coldest alkenone temperatures (∼ 16.0 ∘C) have
been obtained in core MIN2 during the LIAa and the warmest (∼ 18.4 ∘C) in core MR3.3 during the MCA. Values near the average of
maxima SST (from 17.9 to 18.4 ∘C) are observed more frequently
during the TP, RP, and MCA, while temperatures during the onset of MCA and LIA
show many values closer to the average of minima SST (ranging from 16.0 to
16.2 ∘C). Abrupt coolings are observed during the LIA and some
events during MCA (-0.8 ∘C 100 yr-1) and to a lesser extent
during the LIA–IE transition (-0.5 ∘C 100 yr-1). The highest
warming rates are recorded during the MCA (0.4 ∘C 100 yr1)
and also during the RP.
Mg / Ca–SST vs. alkenone–SST records
In this section, the uncertainties of the alkenone, 1.1 ∘C, have
been calculated from the estimated standard error of the calibration (see
Sect. 3.6) and those of Mg / Ca–SST include the analytical precision and
reproducibility and the standard error of the calibration. The measured
Mg / Ca–SST and alkenone–SST averages are identical within error (16.9 ± 1.4 ∘C vs. 17.2 ± 1.1 ∘C), but the temperature
range of the Mg / Ca records shows higher amplitude (see Sects. 5.2 and 5.4).
The similarity in SST averages of both proxies does not reflect the different
habitat depths, since alkenones should mirror the surface photic layer
(< 50 m), with relatively warm SST, while G. bulloides has the capability to
develop in a wider and deeper environment (Bé, 1977; Pujol and
Vergnaud-Grazzini, 1995; Ternois et al., 1996; Sicre et al., 1999; Ganssen
and Kroon, 2000; Schiebel et al., 2002; Rogerson et al., 2004; Thornalley et
al., 2009), where lower SST would be expected.
The enhanced Mg / Ca–SST variability is reflected in the short-term
oscillations, at centennial timescales, which are larger in the Mg / Ca
record with oscillations over 0.5 ∘C. This larger Mg / Ca–SST
variability could be attributed to the highly restricted seasonal character
of the signal, which purely reflects SST changes during the spring season.
However, the coccolith signal integrates a wider time period from autumn to
spring (Rigual-Hernández et al., 2012, 2013) and, consequently, changes
associated with specific seasons become more diluted in the resulting
averaged signal.
The annual mean SST corresponding to a Balearic site is 18.7 ± 1.1 ∘C, according to the integrated values of the upper 50 m
(Ternois et al., 1996; Cacho et al., 1999a) of the GCC-IEO database between
January 1994 and July 2008. Our core-top records represent the last decades
and show SST values closer to the annual mean in the case of alkenone–SST, whereas
the Mg / Ca–SST record shows slightly lower values.
The U37k′–SST records in the western Mediterranean Sea have
been interpreted to represent annual mean SST (i.e. Cacho et al., 1999a;
Martrat et al., 2004) but seasonal variations in alkenone production could
play an important role in the U37k′–SST values
(Rodrigo-Gámiz et al., 2014). Considering that during the summer months
the Mediterranean Sea is a very stratified and oligotrophic sea, reduced
alkenone production during this season could be expected (Ternois et al.,
1996; Sicre et al., 1999; Bárcena et al., 2004; Versteegh et al., 2007;
Hernández-Almeida et al., 2011). This observation is supported by
results from sediment traps located in the GoL showing very low coccolith
fluxes during the summer months (Rigual-Hernández et al., 2013), while
they exhibit higher values during autumn, winter, and spring, reaching
maximum fluxes at the end of the winter season, during SST minima. In
contrast, high fluxes of G. bulloides are almost restricted to the upwelling spring
signal, when coccolith fluxes have already started to decrease
(Rigual-Hernández et al., 2012, 2013). This different growth season can
explain the proxy bias in the SST reconstructions, with more smoothed
alkenone–SST signals.
Both Mg / Ca–SST and U37k′–SST records show consistent
cooling trends of about -0.5 ∘C kyr-1 during the studied
period (2700 years), which is consistent with the recent 2 kyr global
reconstruction (McGregor et al., 2015; see Sect. 5.2). The recorded cooling
since the RP SST maxima (∼ 200 CE) is more pronounced in
the Mg / Ca–SST (-1.7 to -2.0 ∘C kyr-1) than in the alkenone
signal (-1.1 ∘C kyr-1). These coolings are larger than those
estimated in the global reconstruction (McGregor et al., 2015) for the last
1200 years (average anomaly method 1: -0.4 to
-0.5 ∘C kyr-1). It should be noted that the global
reconstruction includes alkenone–SST from MIN cores (data published in Moreno et
al., 2012).
Temperature and isotope anomaly records from Minorca (this study)
and data from other regions. (a) δ18Oc
and δ18Osw
(‰ SMOW) Minorca stacks; (b) alkenone–SST anomaly
Minorca stack; (c) Mg / Ca–SST anomaly Minorca stack; (d) warm and cold phases
and δ18OG.ruber recorded by planktonic foraminifera from
the southern Tyrrhenian composite core, with RCI to RCIV showing
Roman cold periods (Lirer et al., 2014); (e) 30-year averages of the PAGES
2k Network (2013) Europe anomaly temperature reconstruction; (f) Greenland
snow surface temperature (Kobashi et al., 2011); and (g) central Europe summer anomaly temperature reconstruction in central Europe
(Büntgen et al., 2011).
The detailed comparison of the centennial SST variability recorded by both
proxy stacks consistently indicates a puzzling antiphase (Fig. 6b and c).
Although the main trends are consistently parallel in both alkenone and
Mg / Ca proxies (r= 0.5; p value = 0) as observed in other regions,
short-term variability appears to have an opposite character. Statistical
analysis of these differences examined by means of Welch's test indicates
that the null hypothesis (means are equal) can be discarded at the 5 %
error level: tobserved (12.446) > tcritical (1.971).
This a priori unexpected proxy difference outlines the relevance of the
seasonal variability for climate evolution and suggests that extreme winter
coolings were followed by more rapid and intense spring warmings.
Nevertheless, regarding the low amplitude of several of these oscillations,
often close to the proxy error, this observation needs to be supported by further constraints as a solid regional feature.
Discussion
Climate patterns during the last 2.7 kyr
The SST changes in the Minorca region have implications for the surface air
mass temperature and moisture source regions that could influence on air
mass trajectories and ultimately precipitation patterns in the western
Mediterranean region (Millán et al., 2005; Labuhn et al., 2015). Recent
observations have identified SST as a key factor in the development of
torrential rain events in the western Mediterranean Basin (Pastor et al.,
2001), constituting a potential source of mass instability that transits
over these waters (Pastor, 2012). In this context, the combined SST and
δ18Osw records can provide information on the connection
between thermal changes and moisture export from the central-western
Mediterranean Sea during the last 2.7 kyr.
The oldest period recorded in our data is the so-called Talaiotic Period
(TP), which corresponds to the ages of antiquity such as the period of ancient Greece in other
geographic areas. Both Mg / Ca–SST and alkenone–SST records are consistent in
showing a general cooling trend from ∼ 500 BCE and reaching minimum
values by the end of the period (∼ 120 BCE; Fig. 6a–b). Very few
other records are available from this time period, which make comparisons of
these trends at regional scale difficult.
One of the most prominent features in the two SST reconstructions,
particularly in the Mg / Ca–SST stack, is the warm SST that predominated during
the second half of the RP (150–400 CE). The onset of the RP was
relatively cold and a ∼ 2 ∘C warming occurred during the
first part of this period. This SST evolution from colder to warmer
conditions during the RP is consistent with the isotopic record of the Gulf
of Taranto (Taricco et al., 2009) and peat reconstructions from north-west
Spain (Martínez-Cortizas et al., 1999), and to some extent with SST
proxies in the south-eastern Tyrrhenian Sea (Lirer et al., 2014). However, none of
these records indicates that the RP was the warmest period of the last 2 kyr. Other records from higher latitudes such as Greenland (Dahl-Jensen et
al., 1998), and northern Europe (Esper et al., 2014), North Atlantic Ocean (Bond et
al., 2001; Sicre et al., 2008), as well as speleothem records from northern Iberia
(Martín-Chivelet et al., 2011) and even the multiproxy PAGES 2K
reconstruction from Europe, suggest a rather warmer early RP than late RP
and, again, none of these records highlights the Roman times as the warmest
climate period of the last 2 kyr. Consequently, these very warm RP
conditions recorded in the Minorca Mg / Ca–SST stack seem to have a regional
character and suggest that climate evolution during this period followed a
rather heterogeneous thermal response along the European continent and
surrounding marine regions.
Moreover, the observed δ18Osw stack of the RP suggests an
increase in the E–P ratio (Fig. 6a) during this period, which has also
been observed in some nearby regions like the Alps (Holzhauser et al., 2005;
Joerin et al., 2006). In contrast, a lake record from southern Spain
indicates relatively high water levels when the δ18Osw stack
indicates the maximum in E–P ratio (Martín-Puertas et al., 2008). This
information is not necessarily contradictory, since enhanced E–P balance in
the Mediterranean could be balanced out by enhanced precipitation in some of
the regions, but more detailed geographical information is required to
interpret these proxy records from distinct areas.
After the RP, during the whole DMA and until the MCA, the Mg / Ca–SST stack
shows a cooling of ∼1 ∘C (-0.2 ∘C 100 yr-1),
which is 0.3 ∘C in the case of the alkenone–SST stack and the
E–P rate decreases. This trend contrasts with the general warming trend
interpreted from the speleothem records of northern Iberia
(Martín-Chivelet et al., 2011) or the transition towards drier
conditions observed in Alboran records (Nieto-Moreno et al., 2011). However,
SST proxies from the Tyrrhenian Sea show a cooling trend after the second
half of the DMA and the Roman IV cold/dry phase (Lirer et al., 2014) that
can be tentatively correlated with our SST records (Fig. 6). This cooling
phase is also documented in the δ18OG.ruber record of the
Gulf of Taranto (Grauel et al., 2013). These heterogeneities in the signals
from the different proxies and regions illustrate the difficulties in
characterizing the climate variability during these short periods and
reinforce the need for a better geographical coverage of individual proxies.
The medieval period is usually described as a very warm period in numerous
regions in the Northern Hemisphere (Hughes and Diaz, 1994; Mann et al.,
2008; Martín-Chivelet et al., 2011), but this interpretation is
challenged by an increasing number of studies (i.e. Chen et al., 2013). The
Minorca SST stacks also show the occurrence of significant temperature
variability that does not reflect a specific warm period within the last 2 kyr (Fig. 6). An important warming event is observed at
∼ 1000 CE,
followed by a later cooling with minimum values at about 1200 CE
(Fig. 6). Higher temperature variability is found in Greenland records
(Kobashi et al., 2011), while an early warm MCA and posterior cooling is also
observed in temperature reconstructions from central Europe
(Büntgen et al., 2011) and in the European multi-proxy 2 kyr
stack of the PAGES 2K Consortium (2013). Nevertheless, all these proxies
agree in indicating overall warmer temperatures during the MCA than during
the LIA. At the MCA–LIA transition, a progressive cooling and a change in
oscillation frequency before and after the onset of LIA are recorded. This
transition is consistent with the last rapid climate change (RCC) described
in Mayewski et al. (2004).
In the context of the Mediterranean Sea, the lake, marine, and speleothem
records consistently agree in showing drier conditions during the MCA than
during the LIA (Moreno et al., 2012; Chen et al., 2013; Nieto-Moreno et al.,
2013; Wassenburg et al., 2013). Examination of the δ18sw stack
shows several oscillations during the MCA and LIA but no clear
differentiation between these periods can be inferred from this proxy,
indicating that reduced precipitation also involved reduced evaporation in
the basin and that the E–P balance recorded by the δ18Osw proxy was
not modified. The centennial-scale variability found in both the Mg / Ca–SST and δ18Osw stack reveals that higher E–P conditions
existed during the warmer intervals (Fig. 6a and c).
According to the Mg / Ca–SST stack, the LIA stands out as a period of high
thermal variability in which two substages can be differentiated, a first
involving large SST oscillations and warm average temperatures (LIAa) and a
second substage with short oscillations and cold average SST (LIAb). We
suggest that the LIAa interval could be linked to the Wolf and Spörer
solar minima and that the LIAb corresponds to Maunder and Dalton cold
events, in agreement with previous observations (i.e. Vallefuoco et al.,
2012).
These two LIA substages are also present in the Greenland record (Kobashi et
al., 2011). The intense cooling drop (0.8 ∘C 100 yr-1) at the
onset of the LIAb is in agreement with the suggested coolings of 0.5 and
1 ∘C in the Northern Hemisphere (i.e. Matthews and Briffa, 2005;
Mann et al., 2009). These two steps within the LIA are better reflected in
the Mg / Ca–SST stack than in the alkenone–SST stack. This is also the case of
the alkenone records in the Alboran Sea (Nieto-Moreno et al., 2011), which
may result from smaller SST variability of the alkenone proxies (see Sect. 5.5).
In terms of humidity, the LIA represents a period of increased runoff in the
Alboran record (Nieto-Moreno et al., 2011). Available lake level
reconstructions from southern Spain also show progressive increases after the
MCA, reaching maximum values during the LIAb (Martín-Puertas et al.,
2008). Different records of flood events in the Iberia Peninsula also report
a significant increase in extreme events during the LIA (Barriendos and Martin-Vide,
1998; Benito et al., 2003; Moreno et al., 2008). These conditions are
consistent with the described enhanced storm activity over the GoL in this
period (Sabatier et al., 2012), explaining the enhanced humidity transport
towards the Mediterranean Sea as a consequence of the reduced E–P ratio
observed in the δ18Osw, particularly during the LIAb (Fig. 6a).
The end of the LIA and onset of the IE is marked with a warming phase of
about 1 ∘C in the Mg / Ca–SST stack and a lower-intensity change in
the alkenone–SST stack. This initial warm climatic event is also documented
in other Mediterranean regions (Taricco et al., 2009; Marullo et al., 2011;
Lirer et al., 2014) and Europe (PAGES 2K Consortium, 2013), which is
coincident with a total solar irradiance (TSI) enhancement after Dalton
minima. The two Minorca SST stacks show a cooling trend by the end of the
record, which does not seem to be consistent with the instrumental
atmospheric records. In the western Mediterranean, warming has been registered
in two main phases: from the mid-1920s to 1950s and from the mid-1970s
onwards (Lionello et al., 2006). The Minorca stacks do not show this warming,
but they do not cover the second warming period. Nevertheless, the
instrumental data from the beginning of the 20th century in the western
Mediterranean do not display any warming trends before the 1980s
(Vargas-Yáñez et al., 2010).
Temperature and isotope anomaly records from Minorca (this study)
and data from other regions and with external forcings: (a) total solar
irradiance (Steinhilber et al.,
2009, 2012), (b) δ18Osw Minorca stacks, (c) Atlantic
Multidecadal Oscillation (AMO; Gray et al., 2004), (d) North Atlantic Oscillation (NAO) reconstructions (Olsen
et al., 2012; Trouet et al., 2009; and, for the last millennium, Ortega et
al., 2015), (e) Mg / Ca–SST anomaly Minorca stack, (f) summer insolation at 40∘ N (Laskar et al., 2004), (g) alkenone–SST anomaly
Minorca stack, and (h) palaeostorm activity in the Gulf of Lion (Sabatier et
al., 2012).
Climate forcing mechanisms
The general cooling trend observed in both Mg / Ca–SST and alkenone–SST stacks
shows a good correlation with the evolution of summer insolation in the
North Hemisphere, which dominates the present annual insolation balance
(r= 0.2 and 0.8, p value ≤ 0.007, respectively; Fig. 7). In numerous
records from the Northern Hemisphere (i.e. Wright, 1994; Marchal et al., 2002;
Kaufman et al., 2009; Moreno et al., 2012), this external forcing has also
been proposed to control major SST trends during the Holocene period. In
addition, summer insolation seems to have had significant influence in the
decreasing trend of the isotope records during the whole period spanned
(r= 0.4, p value = 0), as has been suggested in, for example, Ausín et al. (2015). In any case, a different forcing mechanism needs to account
for the centennial-scale variability of the records, e.g. increased
volcanism in the last millennium (McGregor et al., 2015), although no
significant correlations have been observed between our records and volcanic
reconstructions (Gao et al., 2008).
Solar variability has frequently been proposed to be a primary driver of the
Holocene millennial-scale variability (i.e. Bond et al., 2001). Several
oscillations observed in the TSI record (Fig. 7a), but the correlations with the Mg / Ca–SST
and alkenone–SST stacks are low, since most of the major TSI drops do not
correspond to SST cold events. However, some correlation is observed between
TSI and alkenone SSTs (r= 0.5, p value = 0). In any case, TSI does not seem
to be the main driver of the centennial-scale SST variability in the studied
records.
One of the major drivers of the Mediterranean interannual variability in
the Mediterranean region is the NAO (Hurrell,
1995; Lionello and Sanna, 2005; Mariotti, 2011). Positive NAO indexes are
characterized by high atmospheric pressure over the Mediterranean Sea and
increases of the E–P balance (Tsimplis and Josey, 2001). During these
positive NAO periods, winds over the Mediterranean tend to be deviated
towards the north, overall salinity increases, and formation of dense deep
water masses is reinforced as the water exchange through the Corsica Channel
is higher and the arrival of northern storm waves decreases (Wallace and
Gutzler, 1981; Tsimplis and Baker, 2000; Lionello and Sanna, 2005). The
effect of NAO on Mediterranean temperatures is more ambiguous. SST changes
during the last decades do not show significant variability with NAO
(Luterbacher et al., 2004; Mariotti, 2011), although some studies suggest an
opposite response between the two basins, with cooling responses in some
eastern basins and warmings in the western basin during positive NAO
conditions (Demirov and Pinardi, 2002; Tsimplis and Rixen, 2002). Although
still controversial, some NAO reconstructions on proxy records are starting to
become available for the period studied (Lehner et al., 2012; Olsen et al., 2012;
Trouet et al., 2012; Ortega et al., 2015). The last millennium is the
best-resolved period, and it allows a direct comparison with our data to
evaluate the potential link to NAO.
The correlations between our Minorca temperature stacks with NAO
reconstructions (Fig. 7) are relatively low in the case of Mg / Ca–SST
(r= 0.3, p value ≤ 0.002) and not significant in the alkenone stack,
indicating that this forcing is probably not the driver of the main trends
in these records, although several uncertainties still exist about the long
NAO reconstructions (Lehner et al., 2012). If detailed analysis is performed
focussing on the more intense negative NAO phases (Fig. 7), they mostly
appear to correlate with cooling phases in the Mg / Ca stack. The frequency of
these negative events is particularly high during the LIA, and mostly during
its second phase (LIAb), when the coldest intervals of our SST stacks were
observed.
δ18Osw Minorca stack (‰
SMOW) during the last millennium (age is expressed in years CE)
plotted with (a) NAO reconstruction (Ortega et al., 2015) and (b) palaeostorm
activity in the Gulf of Lion (Sabatier et al., 2012). Notice that the NAO
axis is on a descending scale. Grey vertical bars represent negative NAO
phases.
When several different proxy last century records of annual resolution,
tested with some model assimilations (Ortega et al., 2015), are compared
with the last NAO reconstruction, the observed correlations with δ18Osw are not statistically significant. However, the Welch's
test results do not allow for the null hypothesis to be discarded. A coherent
pattern of NAO variability with our δ18Osw reconstruction,
with high (low) isotopic values mainly dominating during positive (negative)
NAO phases, can be observed in the last centuries (Fig. 8). This pattern is
consistent with the described E–P increase during high NAO phases described
for the last decades (Tsimplis and Josey, 2001). The SST stacks also suggest
some degree of correlation between warm SST and high NAO values (Fig. 7), but
a more coherent picture is observed when the SST records are compared to the
Atlantic Meridional Oscillation (AMO) reconstruction: warm SST dominated
during high AMO values (Fig. 9). This pattern of salinity changes related to
NAO and SST to AMO has also been described in climate studies encompassing
the last decades (Mariotti, 2011; Guemas et al., 2014) and confirms the
complex but tight response of the Mediterranean to atmospheric and marine
changes over the North Atlantic Ocean.
The pattern of high δ18Osw at dominant positive NAO
corresponds to a reduction in the humidity transport over the Mediterranean
region as a consequence of high atmospheric pressure (Tsimplis and Josey,
2001). Accordingly, several periods of increased/decreased storm activity in
the GoL (Fig. 8; Sabatier et al., 2012) correlate with low/high values in
the δ18Osw, indicating that, during negative NAO
conditions,
northern European storm waves arrived more frequently in the Mediterranean Sea
(Lionello and Sanna, 2005), contributing to the reduction of the E–P
balance (Fig. 8). Our data also indicate that, during these enhanced storm
periods, cold SST conditions dominated in the region as previously suggested
(Sabatier et al., 2012). Nevertheless, not all the NAO oscillations had
identical expression in the compared records, which is coherent with recent
observations indicating that negative NAO phases may correspond to different
atmospheric configuration modes and impact differently over the western
Mediterranean Sea (Sáez de Cámara et al., 2015). Regarding
the lower part of the record, the maximum SST temperatures and δ18Osw recorded during the RP (100–300 CE) may suggest the
occurrence of persistent positive NAO conditions, which would also be
consistent with a high pressure-driven drop in relative sea level as
has been reconstructed in the north-western Mediterranean Sea (southern
France, -40 ± 10 cm; Morhange et al., 2013).
Mg / Ca–SST and alkenone–SST Minorca anomaly stacks during the last
centuries plotted with AMO reconstruction (Gray et al., 2004).
It is interesting to note that during the DMA a pronounced and intense
cooling event is recorded in the Mg / Ca–SST stack at about 500 CE. Several
references document in the scientific literature the occurrence of a dimming of the sun at 536–537 CE (Stothers, 1984). This event,
based on ice core records, has been linked to a tropical volcanic eruption
(Larsen et al., 2008). Tree-ring data reconstructions from Europe and also
historical documents indicate the persistence during several years
(536–550 CE) of what is described as the most severe cooling across the Northern
Hemisphere during the last two millennia (Larsen et al., 2008). Despite the
limitations derived from the resolution of our records, the Mg / Ca–SST stack
record may have caught this cooling, which would prove the robustness of our
age models (see Supplement for age model development).