Introduction
Isotopic compositions of late-glacial precipitation can be preserved in
groundwaters, cave calcite, glacial ice, ground ice and lake sediments.
These records have been used to better understand past climate changes for
more than a half century (e.g., Münnich, 1957; Thatcher et al., 1961;
Münnich et al., 1967; Pearson and White, 1967; Tamers, 1967; Gat et al.,
1969). Each type of isotopic proxy record is distinguished by its temporal
resolution, preservation of one or both 18O / 16O and
2H / 1H ratios, and frequency on land surface. For example,
groundwater records contain both 18O / 16O and 2H / 1H
ratios with widespread global occurrence, but have a coarser temporal
resolution than other paleoclimate proxies (Rozanski, 1985; Edmunds
and Milne, 2001; Edmunds, 2009; Corcho Alvarado et al., 2011;
Jiráková et al., 2011). Speleothem records, by contrast, have high
temporal resolution but usually only report calcite 18O / 16O ratios
(without fluid inclusion 2H / 1H data) and are less common than
groundwater records (e.g., Harmon et al., 1978, 1979). Late-glacial ice
core and ground ice records have high temporal resolution, can be analysed
for 18O / 16O and 2H / 1H ratios, but are rare on non-polar
lands (Dansgaard et al., 1982; Thompson et al., 1989, 1995, 1997, 1998).
Lake sediment records can have a high temporal resolution, can preserve
18O / 16O and 2H / 1H ratios and are available for a
multitude of globally distributed locations (e.g., Edwards and McAndrews, 1989;
Eawag et al., 1992; Menking et al., 1997; Wolfe et al., 2000; Anderson et
al., 2001; Beuning et al., 2002; Sachse et al., 2004; Morley et al., 2005;
Tierney et al., 2008). However, some lake water proxy isotope records may be
impacted by paleo-lake evaporative isotope effects that obscure the primary
meteoric water signal and mask paleo-precipitation isotope compositions
(e.g., lake sediment calcite, diatom silica; Leng and Marshall, 2004).
This study examines speleothem, ice core and groundwater isotope records,
focusing primarily on the groundwater isotope records due to their relative
density in the published literature in comparison to the more limited number
of published speleothem and ice core records (compilations by Pedro et al.,
2011; Stenni et al., 2011; Clark et al., 2012; Shah et al., 2013; Caley et
al., 2014a). There exist roughly twice as many groundwater reconstructions
of late-glacial to late-Holocene precipitation δ18O
shifts (n= 59) as the combined total of speleothem and ice core records
(n= 27; where δ18O =
(18O / 16Osample) / (18O / 16Ostandard mean ocean water-1) × 1000). A
recent global synthesis of paired precipitation-groundwater isotopic data
demonstrated that modern annual precipitation and modern groundwater isotope
compositions follow systematic relationships with some bias toward winter
and wet-season precipitation (Jasechko et al., 2014). Systematic
rainfall-recharge relationships shown by Jasechko et al. (2014) support our
primary assumption in this study that groundwater isotope compositions
closely reflect meteoric water. Because groundwater records can only
identify climate change occurring over thousands of years due to
hydrodynamic dispersion during multi-millennial residence times (e.g.,
Davison and Airey, 1982; Stute and Deak, 1989), we limit the focus of this
study to meteoric water isotope composition changes from the latter half of
the last glacial time period to the late-Holocene. The latter half of the
last glacial period is defined as ∼ 20 000 to ∼ 50 000 years before present, using the end of the last glacial maximum as
the more recent age limit (∼ 20 000 years before present;
Clark et al., 2009) and the maximum age of groundwater that can be
identified by 14C dating as an approximate upper age limit (i.e.,
groundwater ages more recent than ∼ 50 000 years old).
For brevity, we refer herein to the time period representing the latter half
of the last glacial period (∼ 20 000 to ∼ 50 000 years before present) as the late-glacial (e.g., δ18Olate-glacial).
We adopt a definition of the late-Holocene as occurring within the last 5000 years
following Thompson et al. (2006). Other work proposes the late-Holocene
be defined as within the last 4200 years (Walker et al., 2012), which is
consistent with the 5000 years before present definition (Thompson et al.,
2006) within the practical uncertainty of 14C-based groundwater ages
(±∼ 103 years). Further, although precipitation
isotope compositions have varied over the late-Holocene, groundwater
mixing integrates this variability, prohibiting paleoclimate interpretation
at finer temporal resolutions.
Late-glacial to late-Holocene changes in precipitation isotope
compositions provide important insights into conditions and processes of the
past. Perhaps the two best-constrained global-in-scale differences between
the late-glacial and the late-Holocene are changes to oceanic and
atmospheric temperatures (MARGO Members, 2009; Shakun and Carlson, 2010;
Annan and Hargreaves, 2013), and changes to seawater δ18O
(Emiliani, 1955; Dansgaard and Tauber, 1969; Schrag et al., 1996, 2002).
Atmospheric temperatures have increased by a global average of
∼ 4 ∘C since the last glacial maximum, with greatest
warming at the poles and more modest warming at lower latitudes (Fig. 1;
Shakun and Carlson, 2010; Annan and Hargreaves, 2013). Seawater δ18O during the last glacial maximum was 1.0 ± 0.1 ‰ higher than the modern ocean, as constrained by
paleo-ocean water samples collected from pore waters trapped within sea
floor sediments (Schrag et al., 2002).
The change in surface air temperatures from the last glacial
maximum to the preindustrial era (gridded data from Annan and Hargreaves,
2013). (a) Percentile ranges of temperature changes since the last glacial
maximum for 10 degree latitudinal bands. Blue shading marks the
25th–75th percentile range; thin horizontal lines mark the
10th–90th percentile range. The grey band shows the
globally averaged estimate of temperature change since the last glacial
maximum of -4.0 ± 0.8 ∘C. (b) Gridded surface air
temperature anomaly from the last glacial maximum to the preindustrial era
(data from Annan and Hargreaves, 2013).
![]()
Previous studies have proposed many different interpretations of past
changes to precipitation isotope compositions. Records of
paleo-precipitation δ18O have been used as a proxy for regional
land surface and atmospheric temperature (e.g., Rozanski, 1985; Nikolayev
and Mikhalev, 1995; Johnsen et al., 2001; Grasby and Chen, 2005; Akouvi et
al., 2008; Bakari et al., 2012); however, δ18O-based
paleotemperatures can be complicated by past changes to a variety of other
processes controlling precipitation δ18O, including moisture
sources, upwind rainout, transport pathways, moisture recycling and in-cloud
processes (Ciais and Jouzel, 1994; Masson-Delmotte et al., 2005; Sjostrom
and Welker, 2009). Process-based explanations for observed meteoric water
δ18O variations in proxy records include changes to hurricane
intensity (e.g., Plummer, 1993), large-scale atmospheric circulation
(e.g., Rozanski, 1985; Weyhenmeyer et al., 2000; McDermott et al.,
2001; Pausata et al., 2009; Asmerom et al., 2010; Oster et al., 2015),
aridity (e.g., Wagner et al., 2010), monsoon strength (e.g., Denniston et
al., 2000; Lachniet et al., 2004; Liu et al., 2007; Pausata et al., 2011a),
local seawater δ18O (Wood et al., 2003; Feng et al., 2014),
precipitation seasonality (e.g., Fawcett et al., 1997; Werner et al., 2000;
Cruz et al., 2005), moisture provenance (e.g., Sjostrom and Welker, 2009;
Lewis et al., 2010), storm tracks, climate oscillation modes (e.g., North
Atlantic oscillation), moisture recycling (e.g., Winnick et al., 2013, 2014;
Liu et al., 2014a, b) and groundwater flow path architecture (Purdy et
al., 1996; Stewart et al., 2004; Morrissey et al., 2010; Hagedorn, 2015).
While unravelling these mechanisms and delineating the primary and secondary
processes can be rather challenging, the use of climate models in
combination with robust and extensive precipitation isotope data can resolve
many of these complexities with meaningful interpretations and insight.
The objective of this study is to analyse spatial patterns of measured
late-glacial to late-Holocene precipitation δ18O changes
from published groundwater, ground ice, glacial ice and cave calcite
records, and to compare these measurements with output from five
state-of-the-art isotope-enabled general circulation model simulations of
last glacial maximum and pre-industrial or modern climate conditions.
Synthesizing paleowater δ18O records provides an important
constraint for isotope-enabled general circulation model simulations of
atmospheric and hydrologic conditions during glacial climate states (Jouzel
et al., 2000). We combine a new global compilation of late-glacial
groundwater and ground ice isotope data (n= 59) with existing compilations
for speleothems (n= 15; Shah et al., 2013) and ice cores (n= 12; Pedro et
al., 2011; Stenni et al., 2011; Clark et al., 2012; Caley et al., 2014a).
This compilation of late-glacial groundwater isotope compositions builds
from earlier reviews of European and African paleowater isotope compositions
(Rozanski, 1985; Edmunds and Milne, 2001; Darling, 2004; Edmunds, 2009;
Négrel and Petelet-Giraud, 2011; Jiráková et al., 2011).
Data set and methods
In order to examine spatial patterns of change to meteoric water δ18O values we compiled δ18O, δ2H, δ13C and 14C data from 1713 groundwater samples
collected from 59 aquifer systems reported in 76 publications (data and
primary references presented in the Supplement). δ13C, 3H
and 14C data were used to estimate groundwater age (details within
Supplement). Changes to precipitation δ18O values over time
were determined by comparing groundwater isotope compositions of the
late-Holocene (δ18Olate-Holocene defined here as
less than 5000 years before present; Thompson et al., 2006) and the latter
half of the last glacial time period (δ18Olate-glacial:
20 000 to ∼ 50 000 years before present). We acknowledge that
these two relatively long time intervals – necessarily long in order to
examine groundwater isotope records – integrate precipitation δ18O variability over the course of each time interval. The
late-Holocene time interval integrates known precipitation δ18O variability (e.g., Aichner et al., 2015), and the late-glacial
time interval likely incorporates groundwater preceding the last glacial
maximum, potentially during Marine Isotope Stage 3 or even older glacial
time periods due to large uncertainties in 14C-based groundwater ages
(Supplement).
Proxy-based meteoric water δ18O changes from the latter half of
the last glacial time period to the late-Holocene are described herein as
measured Δ18Olate-glacial, where measured Δ18Olate-glacial=δ18Olate-glacial-δ18Olate-Holocene. A minimum groundwater age of
20 000 years before present was used to define the late-glacial to remain
consistent with the timing of the last glacial maximum (∼ 20 000 years before present; Clark et al., 2009). Samples having a deuterium
excess of less than zero (deuterium excess =δ2H -
8 × δ18O; Dansgaard, 1964) and falling along
regionally characteristic evaporation δ2H/δ18O
slopes (Gibson et al., 2008) were removed from the analysis to avoid
including groundwater samples impacted by partial evaporation. Further,
studies reporting saltwater intrusion were avoided on the basis of
groundwater δ18O and salinities showing evidence of seawater
mixing (e.g., Schiavo et al., 2009; Yechieli et al., 2009; Hamouda et al.,
2011; Han et al., 2011; Wang and Jiao, 2012; Currell et al., 2013). The 59
compiled groundwater measured Δ18Olate-glacial values
are unevenly distributed among western Europe (n= 10), eastern Europe and
the Middle-East (n= 12), Africa (n= 17), southeastern Asia (n= 6),
Australia, Oceania and the Malay Archipelago (n= 2), South America (n= 2),
temperate and subtropical North America (n= 8) and the High Arctic (n= 2).
Half of the compiled groundwater records are located in the tropics or
subtropics (that is, within 35∘ of the equator; n= 29) and half
are located in the extra-tropics (n= 30).
Speleothem and ice core isotope proxy records were also compiled. Lacustrine
sediment δ18O records are not considered in this study because
these records may preserve meteoric waters impacted by evaporative isotope
effects (Leng and Marshall, 2004). Speleothem and ice core measured Δ18Olate-glacial values were calculated by subtracting average
δ18O values for each of the two time intervals defined for the
groundwater records: the late-Holocene (< 5000 years before
present) and latter half of the last glacial time period (20 000 to 50 000 years before present). This step effectively lowered the temporal resolution
of speleothem and ice core precipitation isotope records to be consistent
with the temporal resolution of the groundwater records. A correction factor
was applied to speleothem δ18O values to account for different
H2O-CaCO3 isotopic fractionation factors during the
late-glacial and the late-Holocene because of differing land surface
temperatures during each time period (details presented within Supplement).
Simulated Δ18Olate-glacial values were compiled from
five isotope-enabled general circulation models (simulated Δ18Olate-glacial=δ18Olast glacial maximum-δ18Opre-industrial): CAM3iso (e.g., Noone and Sturm,
2010; Pausata et al., 2011a), ECHAM5-wiso (e.g., Werner et al., 2011),
GISSE2-R (e.g., Schmidt et al., 2014; LeGrande and Schmidt, 2008, 2009),
IsoGSM (e.g., Yoshimura et al., 2003) and LMDZ4 (e.g., Risi et al., 2010a).
ECHAM5-wiso and IsoGSM outputs are for modern climate rather than
pre-industrial conditions; however, the difference between the isotopic
composition of pre-industrial and modern climate are expectedly small
compared to late-glacial to late-Holocene δ18O shifts. An
offset factor was applied to simulated mean seawater δ18O in
all five models (Table S1 in the Supplement) to account for known glacial-interglacial changes
to seawater δ18O (Emiliani, 1955; Dansgaard and Tauber, 1969;
Schrag et al., 1996, 2002). Possible spatial differences in seawater δ18O changes from the last glacial maximum to the pre-industrial time
period are not incorporated into simulations with prescribed sea surface
temperatures (CAM3iso, ECHAM5-wiso, IsoGSM, LMDZ4) but are simulated by the
coupled ocean-atmosphere simulation of GISSE2-R (Table S1).
GISSE2-R was submitted to the CMIP5 archive and participated in PMIP3. LMDZ4
was submitted to the CMIP3 archive. ECHAM5 and CAM3iso did not participate
in CMIP5, while IsoGSM uses different boundary conditions than proposed for
CMIP5 (Yoshimura et al., 2008). The five models span a range of
spatio-temporal resolutions and isotopic/atmospheric parameterizations
described in detail in the above references. A selection of the inter-model
similarities and differences are summarized in Table S1.
For clarity, empirical Δ18Olate-glacial values that are
based on measured isotope contents of groundwater, speleothem, ground ice or
ice core records are referred to herein as measured Δ18Olate-glacial; simulated precipitation isotope compositions
obtained from general circulation model results are referred to as
simulated Δ18Olate-glacial. We acknowledge that the general
circulation models explicitly analyse the last glacial maximum and the
pre-industrial climate conditions (i.e., simulated Δ18Olate-glacial=δ18Olast glacial maximum-δ18Opre-industrial), whereas proxy record reconstructions of
Δ18Olate-glacial integrate hydroclimatology over
multi-millennial timescales that are different from the model simulations.
Results and discussion
Measured <inline-formula><mml:math display="inline"><mml:mrow><mml:msup><mml:mi mathvariant="normal">Δ</mml:mi><mml:mn>18</mml:mn></mml:msup></mml:mrow></mml:math></inline-formula>O<inline-formula><mml:math display="inline"><mml:msub><mml:mi/><mml:mtext>late-glacial</mml:mtext></mml:msub></mml:math></inline-formula> values
Measured groundwater (n= 59), speleothem (n= 15) and ice core (n= 12)
Δ18Olate-glacial values are presented in Fig. 2
(references presented in the Supplement). Measured Δ18Olate-glacial values range from -7.1 ‰
(i.e., δ18Olate-glacial < δ18Olate-Holocene) to +1.7 ‰ (i.e.,
δ18Olate-glacial > δ18Olate-Holocene). Three-quarters of the compiled records have
negative measured Δ18Olate-glacial values and
one-quarter of compiled records have positive measured Δ18Olate-glacial values. Most groundwater-based late-glacial to
late-Holocene shifts fall along δ2H/δ18O slopes of
∼ 8 (Fig. S58 in the Supplement), suggesting that most groundwaters record
temporal shifts to precipitation isotope contents rather than to soil
evaporation isotope effects (see Evaristo et al., 2015). More than 80 % of
records with positive measured Δ18Olate-glacial values
are located within 35∘ of the equator and within 400 km of the
nearest coastline (e.g., Bangladesh Δ18Olate-glacial of
+1.5 ‰, less than 300 km from the coast; Figs. 2–4). In comparison, negative measured Δ18Olate-glacial
values are found in both coastal regions and farther inland. Negative
measured Δ18Olate-glacial values of the greatest
magnitude are located at high latitudes (e.g., northwestern Canada, latitude
64∘ N: Δ18Olate-glacial of
-5.5 ‰; northern Russia latitude 72∘ N:
-5.4 ‰) and far from coastlines (e.g., Hungary:
-3.7 ‰, ∼ 500 km from Atlantic Ocean;
Peru: -6.3 ‰, ∼ 2000 km from Atlantic
Ocean, the modern moisture source to Peru; Garreaud et al., 2009). Greenland
and Antarctic ice cores have negative measured Δ18Olate-glacial values that are of greater magnitude than
non-polar measured Δ18Olate-glacial values (Antarctic
and Greenland Δ18Olate-glacial values range from
-3.6 to -7.1 ‰; Fig. 3).
Meteoric water δ18O change from the
late-glacial (20 000 to ∼ 50 000 years ago) to the
late-Holocene (within past ∼ 5000 years; average Δ18Olate-glacial values shown, where Δ18Olate-glacial=δ18Olate-glacial-δ18Olate-Holocene). The low temporal resolution of
groundwater records means that δ18O variations within each time
period are smoothed and likely represent unequal temporal weighting.
References for measured meteoric water δ18O changes for ice
cores, groundwater and cave calcite are presented in the Supplement.
![]()
Our synthesis shows that measured Δ18Olate-glacial
values in the tropics are closer to 0 ‰ (i.e., no
change) than Δ18Olate-glacial values at high latitudes
and continental interiors that generally have high magnitude, negative
Δ18Olate-glacial values. High magnitude, negative
measured Δ18Olate-glacial values are most common where
present-day precipitation δ18O values are at a minimum (e.g.,
Bowen and Wilkinson, 2002). This broad spatial pattern is consistent with
the non-linear isotopic distillation of air masses undergoing progressive
rainout (i.e., Rayleigh distillation). Because seawater δ18O
values were ∼ 1 ‰ higher-than-modern
during the last glacial maximum (Schrag et al., 1996, 2002), our finding
that the majority of measured Δ18Olate-glacial values
are negative suggests that isotopic distillation of air masses was greater
during the late-glacial than under present climate. This finding is
consistent with land surface temperature reconstructions that show larger
glacial-to-modern changes to land temperatures at high latitude and
continental settings (Fig. 1; Annan and Hargreaves, 2013). Tropical versus
extratropical patterns of late-glacial/late-Holocene temperature
change (Fig. 1a) are broadly similar to measured Δ18Olate-glacial values (Fig. 3), where both temperature and
isotope shifts are greater at high latitudes relative to the equator.
Therefore, it is possible that the larger late-glacial to
late-Holocene temperature shifts at the poles relative to the equator may
have served to amplify the non-linear, Rayleigh relationship describing the
heavy isotope depletion of air masses undergoing progressive rainout during
transport from lower to higher latitudes. Further, the late-glacial was
characterized by (i) lower-than-modern atmospheric temperatures with larger
coastal-inland gradients, and (ii) lower-than-modern eustatic sea level
leading to longer overland atmospheric transport distances. Each of these
late-glacial/late-Holocene changes favours stronger-than-modern
isotopic distillation of air masses transported inland from the coast during
the late-glacial (Dansgaard, 1964; Rozanski, 1993; Winnick et al.,
2014), potentially contributing to the broad, global observation that most
(77 %) δ18Olate-Holocene values exceed δ18Olate-glacial values on continents.
Latitudinal variations of Δ18Olate-glacial
values of groundwater (circles, each circle is one aquifer), ice cores
(diamonds) and cave calcite (i.e., triangles; where Δ18Olate-glacial=δ18Olate-glacial-
δ18Olate-Holocene). Dashed lines mark 10∘
zonal mean simulated Δ18Olate-glacial values from five
different general circulation models: CAM3iso, ECHAM5-wiso, GISSE2-R, IsoGSM
and LMDZ4 (Yoshimura et al., 2003; Legrande and Schmidt, 2008, 2009; Risi et
al., 2010a; Noone and Sturm, 2010; Pausata et al., 2011a; Werner et al.,
2011).
![]()
Pairings of groundwater and speleothem records are available within
∼ 500 km of one another in the southwestern USA, central China
and Israel. Southwestern USA speleothem and groundwater records
∼ 400 km apart show similar Δ18Olate-glacial values, with San Juan Basin groundwaters
having a measured Δ18Olate-glacial value of
-2.5 ± 1.0 ‰ (Phillips et al., 1986) and
speleothems ∼ 400 km to the south having measured Δ18Olate-glacial values of -3.0 ± 1.2 and -3.4 ± 0.4 (Asmerom et al., 2010; Wagner et al., 2010). Central China speleothem
and groundwater records ∼ 200 km apart overlap within
uncertainty margins (i.e., Δ18Olate-glacial values of
-1.1 ± 1.7 and +0.3 ± 2.1 ‰; Cai et al., 2010). Israeli speleothem and
groundwater records ∼ 100 km apart have different measured
Δ18Olate-glacial values. Two Israeli groundwater
Δ18Olate-glacial records were compiled; the coastal
Israeli aquifer has a Δ18Olate-glacial value of
+0.3 ± 0.4 ‰ (Yechieli et al., 2009), whereas
groundwater of the Dead Sea Rift Valley has a Δ18Olate-glacial value of -1.8 ± 0.6 ‰ (Burg et al., 2013). Speleothem records have
Δ18Olate-glacial values close to
+1 ‰ (Frumkin et al., 1999; Bar-Matthews et al.,
2003). In northern Turkey, speleothem and groundwater separated by
∼ 150 km have measured Δ18Olate-glacial
values that differ by ∼ 3 ‰ (speleothem
Δ18Olate-glacial-5.7 ± 0.4 ‰ versus groundwater Δ18Olate-glacial of -2.8 ± 1.0 ‰;
Fleitmann et al., 2009; Arslan et al., 2013, 2015). While the locations of
the groundwater and speleothem records differ, the compiled data suggest
that groundwater and speleothem Δ18Olate-glacial values
may capture different Δ18Olate-glacial values under
similar climate conditions.
A number of potential processes could bias the preservation of precipitation
isotope composition in ice core, speleothem or groundwater archives (Wang et
al., 2001; Thompson et al., 2006; Edmunds, 2009). For example, groundwater
and speleothem archives preserve only the isotope record of precipitation
that traverses the vadose zone. Recent global analyses of paired
precipitation-groundwater isotope compositions show that winter
(extratropics) and wet season (tropics) precipitation contributes
disproportionately to recharge (Jasechko et al., 2014), meaning that
paleoclimate records may be more sensitive to changes to winter and wet
seasons than summer or dry season (Vogel et al., 1963; Simpson et al., 1972;
Grabczak et al., 1984; Harrington et al., 2002; Jones et al., 2000; Darling,
2004; Partin et al., 2012). Similarly, groundwater isotope records are
unlikely to represent constant and continuous recharge fluxes during the
late-Holocene or the late-glacial (McIntosh et al., 2012). Modern
groundwater recharge fluxes are highest in humid climates (Wada et al.,
2010). Groundwater δ18O records only represent precipitation
that recharges aquifers, meaning that groundwater-based Δ18Olate-glacial values could be biased to subintervals (e.g.,
interstadials, pluvial periods) within the late-Holocene and late-glacial
intervals when recharge fluxes were at local maxima. Speleothem records may
be further complicated by processes impacting the timing of calcite
precipitation. Recent modelling suggests that calcite precipitation in caves
located outside of the tropics is greatest during the cool season and
reduced during summer months due to changes in ventilation, meaning that
higher latitude speleothems record oxygen isotope compositions biased to
cool season climate change (James et al., 2015). Other recent work suggests
that speleothem δ18O data may be impacted by disequilibrium
isotope effects (Asrat et al., 2008; Daëron et al., 2011; Kluge and
Affek, 2012; Kluge et al., 2013) or by partial evaporation of drip waters
resulting in 18O-enrichment (e.g., Cuthbert et al., 2014a) and greater
fractionation due to evaporative cooling (Cuthbert et al., 2014b),
potentially obscuring the preservation of primary precipitation isotope
contents in the speleothem record. Compiled ice core records may have been
influenced by post-depositional exchanges of ice with atmospheric vapour
(Steen-Larsen et al., 2014). The impact of atmospheric vapour exchanges on
ice core isotope records remains poorly understood. Potential biases in the
preservation of precipitation δ18O differ among groundwater,
glacial ice, and speleothem records, meaning that co-located records of
differing record-type may preserve different Δ18Olate-glacial values under similar climate conditions.
Finally, all proxy records may be impacted by past changes in the
seasonality of precipitation, which can substantially impact annual
precipitation δ18O values (e.g., Werner et al., 2000).
We cannot rule out the possibility that changes in seasonal biases of proxy
record preservation occurred between the late-glacial and the
late-Holocene and have impacted measured Δ18Olate-glacial values. Further, the chronologies of
groundwaters and ice core records have uncertainties on the order of
thousands of years, meaning that the time intervals used to calculate
measured Δ18Olate-glacial values may be inaccurate.
However, the plateauing of isotope content observed in most regional
aquifers for 0–5000 years before present and for > 20 000 years
before present supports our interpreting these data as records of
late-glacial to late-Holocene isotopic shifts (see figures in the
Supplement). Notwithstanding potential δ18O preservation biases
and chronology uncertainties, the global data synthesized here show patterns
consistent with the enhanced distillation of advected air masses originating
as (sub)tropical ocean evaporate and undergoing progressive, poleward
rainout under cooler-than-modern late-glacial temperatures.
Measured Δ18Olate-glacial value variability
with distance to the nearest coast (Δ18Olate-glacial
=δ18Olate-glacial-δ18Olate-Holocene). Tropical and subtropical locations are
shown in deep blue (< 35∘ absolute latitude),
extra-tropical sites are shown in light grey (> 35∘
absolute latitude). The shape of each point corresponds to groundwater and
ground ice (circles) or cave calcite (i.e., speleothems; triangles). Error
bars mark one standard deviation from the mean.
![]()
Simulated <inline-formula><mml:math display="inline"><mml:mrow><mml:msup><mml:mi mathvariant="normal">Δ</mml:mi><mml:mn>18</mml:mn></mml:msup></mml:mrow></mml:math></inline-formula>O<inline-formula><mml:math display="inline"><mml:msub><mml:mi/><mml:mtext>late-glacial</mml:mtext></mml:msub></mml:math></inline-formula> values
Simulated precipitation Δ18Olate-glacial values from
five general circulation models are presented in Fig. 5. At least four of
the five models agree on the sign of simulated Δ18Olate-glacial values – that is values consistently above or
consistently below zero – for 68.8 % of grid cells covering Earth's
surface (68.7 % of over-ocean areas and 68.9 % of over land areas;
multi-model calculation completed using three of four models as a threshold at
high-latitudes where IsoGSM data were unavailable). Simulated Δ18Olate-glacial values are consistently negative over the
North Atlantic Ocean and the Fennoscandian and Laurentide ice sheets and
consistently positive over most of the tropical oceans, whereas poorer
agreement is found over tropical land surfaces. The negative simulated
Δ18Olate-glacial values over the Northern Hemisphere
ice sheets and North Atlantic are likely driven by the difference in ice
sheet topography and sea ice cover, between the late-glacial and
pre-industrial climate. The late-glacial to late-Holocene change in ice
sheet topography and sea ice cover impacted surface temperatures, which were
more than ∼ 20 ∘C cooler over most of present-day
Canada during the last glacial maximum (Fig. 1). Cooler temperatures in
conjunction with ice sheet topography (> 3000 m elevations; e.g.,
Peltier, 1994) enhanced Rayleigh distillation for air masses transecting
Northern Hemisphere ice sheets, as evidenced by systematically low measured
and simulated δ18Olate-glacial values in these regions
(Figs. 2, 3 and 5).
Simulated precipitation δ18O differences between the
last glacial maximum and pre-industrial time periods (i.e., δ18Olast glacial maximum – δ18Opre-industrial) from five general circulation models: CAM3iso,
ECHAM5-wiso, GISSE2-R, IsoGSM and LMDZ4 (Yoshimura et al., 2003; Legrande
and Schmidt, 2008, 2009; Risi et al., 2010a; Noone and Sturm, 2010; Pausata
et al., 2011a; Werner et al., 2011). Circles (groundwater), triangles
(speleothems) and diamonds (ice cores) show measured Δ18Olate-glacial values from paleoclimate proxy records (Fig. 1, original data presented in Tables S2–S5). The panel entitled
“Composite” shows the multi-model ensemble median simulated Δ18Olate-glacial value where at least four of the five models
agree on the sign of simulated Δ18Olate-glacial values
(i.e., positive or negative; all five model simulations of δ18Olast glacial maximum – δ18Opre-industrial were used to calculate multi-model median shown
in “Composite”).
![]()
A comparison of simulated Δ18Olate-glacial values over
tropical Africa, South America and Oceania shows inter-model disagreement
(Fig. 5). Different tropical simulated Δ18Olate-glacial values among the models reflect the different
isotopic parameterizations, inter-model spread in simulated precipitation
rates, and seawater δ18O specifications used in each model
(Supplement). Inter-model spread in simulated Δ18Olate-glacial values in some regions highlights the
importance of this global synthesis of measured Δ18Olate-glacial values as a constraint for isotope-enabled
climate simulations. Another potential source for the model disagreement is
introduced by the different ice-sheet topography used in each model.
CAM3Iso, IsoGSM and LMDZ4 used Ice 5G (Peltier, 1994) as advised for PMIP2
(Braconnot et al., 2007), whereas the GISSE2 replaces Ice 5G Laurentide ice
with that of Licciardi et al. (1999) and ECHAM5-wiso uses ice topography
from PMIP3 (Braconnot et al., 2007, 2012; PMIP3 follows ice sheet topography
blended from multiple ice sheet reconstructions: Argus and Peltier, 2010;
Toscano et al., 2011). Ice sheet topography is an important driver of
simulated temperature, precipitation and atmospheric circulation during the
last glacial maximum (e.g., Justino et al., 2005; Pausata et al., 2011b;
Ullman et al., 2014). Therefore, it is likely that inter-model differences
in paleo-ice sheet topographies impacts atmospheric circulation and thus
high latitude simulated Δ18Olate-glacial values
reported in this study (Fig. 5).
Differences in the specification of initial seawater δ18O may
also lead to inter-model differences in simulated Δ18Olate-glacial values. Seawater δ18O is set to
be globally homogenous in CAM3Iso, IsoGSM and LMDZ4, and heterogeneous in
ECHAM5-wiso (using modern gridded seawater δ18O heterogeneity
of LeGrande and Schmidt, 2006) and GISSE2-R (coupled atmosphere-ocean model;
seawater δ18O is calculated by the ocean model). Including
surface ocean δ18O heterogeneities in model simulations impacts
land precipitation δ18O by up to ∼ 1.5 ‰ relative to simulations with homogenous seawater
δ18O (LeGrande and Schmidt, 2006). However, different seawater
δ18O specifications cannot account for all inter-model
differences in simulated Δ18Olate-glacial values.
The models also show deficiencies in simulating measured Δ18Olate-glacial values in the tropics, particularly over
tropical Africa. This finding could, in part, relate to the high sensitivity
of precipitation δ18O to convective parameterizations (Lee et
al., 2009; Field et al., 2014), although future research is required to test
this. Another reason may be that the measured Δ18Olate-glacial integrates the hydroclimatological signal
over multi-millennial timescales, whereas the simulated Δ18Olate-glacial values explicitly explore last
glacial maximum and pre-industrial/present-day climate conditions. The
smeared temporal resolution of groundwater-based measured Δ18Olate-glacial values due to storage and mixing in the
aquifer precludes an ideal comparison of measured versus simulated Δ18Olate-glacial values. Further, as previously discussed in
Sect. 3.1, the measured Δ18Olate-glacial values are
susceptible to a number of potential biases that may obscure the magnitude
and direction of late-glacial to late-Holocene precipitation δ18O changes. Notwithstanding, models correctly simulate the sign of
measured Δ18Olate-glacial values (i.e., positive or
negative) in the extratropics more frequently than in the tropics. Better
agreement in the sign of simulated versus measured Δ18Olate-glacial values in the extra-tropics compared to the
tropics is likely linked to the substantial changes to extra-tropical
ice-sheet topography and sea-ice cover between the two climate states in
northern North America and Europe. Substantial changes to Northern
Hemisphere ice volumes between the late-glacial and the late-Holocene
likely enhanced upwind distillation of air masses leading to high-magnitude,
negative Δ18Olate-glacial values that are well captured
by the climate simulations. However, simulated Δ18Olate-glacial values over Antarctica and Greenland show
large inter-model spread, suggesting that model-based interpretations of
polar ice core records may vary widely among different atmospheric models.
Regional measured and simulated <inline-formula><mml:math display="inline"><mml:mrow><mml:msup><mml:mi mathvariant="normal">Δ</mml:mi><mml:mn>18</mml:mn></mml:msup></mml:mrow></mml:math></inline-formula>O<inline-formula><mml:math display="inline"><mml:msub><mml:mi/><mml:mtext>late-glacial</mml:mtext></mml:msub></mml:math></inline-formula>
values
Australia and Oceania
Measured Δ18Olate-glacial values from Australia and
Oceania fall between -1 and
+1 ‰ (Fig. 2). Australian climate during the last
glacial time period was more arid (Nanson et al., 1992), dustier (Chen et
al., 1993) and cooler (Miller et al., 1997) than present day. Simulated
Δ18Olate-glacial values across Australia are variable
among the five models. Measured Δ18Olate-glacial values
across Oceania have been attributed to temporal changes in the strength of
monsoons and convective rains (Aggarwal et al., 2004; Partin et al., 2007;
Williams et al., 2010) potentially impacted by late-glacial to late-Holocene
shifts in the position of the intertropical convergence zone (Lewis et al.,
2010, 2011).
Southeast Asia
Measured Δ18Olate-glacial values from southeastern Asia
range from -2.3 to +1.7 ‰ .
The highest regional measured Δ18Olate-glacial values
are found in Bangladesh (measured Δ18Olate-glacial of
+1.5 ± 1.3 ‰; Aggarwal et al., 2000) and in
central and southeastern China (measured Δ18Olate-glacial of +0.3 to
+1.7 ‰; Wang et al., 2001; Yuan et al., 2004; Dykoski
et al., 2005; Cai et al., 2010; Yang et al., 2010). General circulation
models have positive simulated Δ18Olate-glacial values
near to the Chinese coasts, but are more variable across western and
northern China (Fig. 5). Chinese speleothem records show near-zero or
positive measured Δ18Olate-glacial values interpreted
to reflect the reduced strength of the East Asian (Wang et al., 2001;
Dykoski et al., 2005; Cosford et al., 2008) or Indian monsoons (Pausata et
al., 2011a). Further research suggests that Chinese speleothem δ18O variations reflect changes to regional moisture sources and the
intensity or provenance of atmospheric transport pathways (LeGrande and
Schmidt, 2009; Dayem et al., 2010; Lewis et al., 2010; Maher and Thompson,
2012; Caley et al., 2014b; Tan, 2014).
North China Plain groundwaters have high-magnitude, negative Δ18Olate-glacial values (measured Δ18Olate-glacial of -2.3 ± 0.6 ‰;
Chen et al., 2003) compared to coastal, more southerly counterparts.
Combining the negative measured Δ18Olate-glacial in
northern China (Chen et al., 2003; Ma et al., 2008; Currell et al., 2012; Li
et al., 2015) with the positive measured Δ18Olate-glacial values in central and southeastern China (Wang
et al., 2001; Yuan et al., 2004; Dykoski et al., 2005; Cai et al., 2010;
Yang et al., 2010) reveals a south-to-north decrease from positive (south)
to negative (north) measured Δ18Olate-glacial values
(Figs. 2 and 6). Previous studies of modern precipitation have identified
increasing precipitation δ18O values from the coast to inland
China during the wet season, sharply contrasting spatial patterns expected
from Rayleigh distillation (Aragúas-Aragúas et al., 1998). A more
recent work suggests that low wet-season precipitation δ18O
values over southern China are controlled by the deflection of westerlies
around the Tibetan Plateau, whereas precipitation δ18O values
over northern China are controlled by local-scale rainfall and below-cloud
raindrop evaporation (Lee et al., 2012). Therefore, measured Δ18Olate-glacial values from southern China may reflect changes
to atmospheric circulation at broader spatial scales, whereas measured
Δ18Olate-glacial values from northern China may
indicate changes to more localized atmospheric conditions impacting
processes such as raindrop evaporation in addition to meso- and
synoptic-scale circulation changes.
Regional proxy record Δ18Olate-glacial values for (a) southeastern Asia, (b) Africa,
(c) Europe, and (d) the contiguous United States of America (where Δ18Olate-glacial=δ18Olate-glacial-
δ18Olate-Holocene). The multi-model ensemble median
simulated Δ18Olate-glacial value is shown as a grid
(0.5 degree smoothing). Groundwater records are represented by circles,
speleothems by triangles, and ice cores by diamonds, labels show measured
Δ18Olate-glacial values for each individual record.
![]()
Africa
Measured Δ18Olate-glacial values from Africa range from
-2.9 to +0.1 ‰ (Figs. 2 and 6). Sixteen of 17 measured Δ18Olate-glacial values from
Africa are negative. Near-zero measured Δ18Olate-glacial values are generally found near to coasts
(e.g., Senegal Δ18Olate-glacial of +0.1 ± 0.8 ‰; Madioune et al., 2014), whereas higher
magnitude, negative measured Δ18Olate-glacial values in
Africa are found farther inland (e.g., Niger Δ18Olate-glacial values of -2.3 ± 2.0 and -2.9 ± 0.9 ‰:
∼ 800 km from the Atlantic coast). General circulation model
Δ18Olate-glacial values show poor agreement with
measured Δ18Olate-glacial over tropical Africa compared
to model-measured comparisons for Europe and North America (Fig. 5), with
positive simulated Δ18Olate-glacial values predicted
over large parts of Africa where negative Δ18Olate-glacial values are measured. Figure 5 shows that
Africa has the largest inter-model and model-measurement disagreements in
the sign of Δ18Olate-glacial values of the continents.
Northern African hydrological processes are influenced by interlinked
controls such as meridional shifts in the position of the intertropical
convergence zone (Arbuszewski et al., 2013) and the strength of Atlantic
meridional overturning circulation (Mulitza et al., 2008). Paleowater
chemistry indicates that northern Africa was at least 2 ∘C cooler
than today (Guendouz et al., 1998) and that westerly moisture transport was
stronger than the present during the late-glacial (Sultan et al., 1997;
Abouelmagd et al., 2012).
Tropical Africa was 2 to 4 ∘C cooler and more arid
than present day at the last glacial maximum (Powers et al., 2005; Tierney
et al., 2008). Early- and late-Holocene rainfall and isotope compositions
were highly variable across Africa (Tierney et al., 2008, 2013; Schefuß et
al., 2011; Otto-Bliesner et al., 2014). Tropical
African rainfall originates from both Indian and Atlantic sources, with
Atlantic-sourced moisture travelling across the Congo rainforest (Levin et
al., 2009). Lower-than-modern continental moisture recycling during the
late-glacial may partially explain negative measured Δ18Olate-glacial values across some regions of inland tropical
Africa (e.g., Risi et al., 2013). Negative measured Δ18Olate-glacial values in tropical Africa could also be
interpreted to reflect higher-than-modern upwind rainout during the
late-glacial (see Risi et al., 2008, 2010b; Lee et al., 2009; Scholl et
al., 2009; Lekshmy et al., 2014; Samuels-Crow et al., 2014); however, this
explanation necessitates stronger-than-modern convection during the
late-glacial, an explanation that would contradict the established
cooler-than-modern land surface temperatures. Therefore, changes to
atmospheric transport distances and vapour origins are more likely
responsible for negative measured Δ18Olate-glacial values across tropical Africa (Lewis et al., 2010).
Europe and the Mediterranean
Measured Δ18Olate-glacial values across Europe, the
Middle-East and the eastern Mediterranean range from
-5.7 to +1.3 ‰. Eighty percent of
measured Δ18Olate-glacial values across these regions
are negative. All five general circulation models agree on negative
simulated Δ18Olate-glacial values across Europe,
consistent with the negative measured Δ18Olate-glacial
values across the majority of Europe. Measured Δ18Olate-glacial values are generally higher in western Europe
(0.0 to -1.0 ‰ in Portugal, the
United Kingdom and France) than in eastern Europe
(-1.0 to -5.7 ‰ in Poland,
Hungary and Turkey; Stute and Deak, 1989; Le Gal La Salle et al., 1996;
Darling et al., 1997; Barbecot et al., 2000; Zuber et al., 2004; Galego
Fernandes and Carreira, 2008; Celle-Jeanton et al., 2009; Varsányi et
al., 2011; Samborska et al., 2013; Arslan et al., 2013). This spatial
pattern of Δ18Olate-glacial values is consistent with
enhanced isotopic distillation of westerlies during the late-glacial due to
cooler-than-modern final condensation temperatures.
High magnitude, negative measured Δ18Olate-glacial
values are located in Turkey and Georgia south and east of the Black Sea
(-2.8 ± 1.0 to -5.7 ± 0.4 ‰; Fleitmann et
al., 2009; Arslan et al., 2013; Melikadze et al., 2014). Westerly air mass
trajectories distal to the Fennoscandian ice sheet topography may not have
changed considerably since the late-glacial over western and central
Europe (Rozanski, 1985; Loosli et al., 2001). Therefore, higher, near-zero
measured Δ18Olate-glacial values in western Europe and
lower, negative measured Δ18Olate-glacial values in
eastern Europe indicate enhanced distillation of advected air masses during
the late-glacial relative to the late-Holocene.
Changes to freeze-thaw conditions of the ground surface between the latter
half of the last glacial time period and the modern climates may have
impacted the seasonality of the fraction of precipitation recharging
aquifers and thus Δ18Olate-glacial (Darling, 2004,
2011; Jasechko et al., 2014). Geomorphic evidence suggests
permafrost covered portions of Hungary at the last glacial maximum,
suggesting that land temperatures may have been up to 15 ∘C cooler
than present day (Fábián et al., 2014), a larger late-glacial to
late-Holocene temperature shift than earlier, noble gas-based
reconstructions (5–7 ∘C; Deák et al., 1987). European pollen
records indicate that northern Europe was tundra-like and that southern
Europe was semi-arid during the last glacial maximum (Harrison and Prentice,
2003; Clark et al., 2012). The European late-glacial to late-Holocene
transition from semi-arid deserts to temperate forests could have lowered
Δ18Olate-glacial values as groundwater recharge ratios
transitioned from more extreme winter-biased (e.g., semi-arid lands during
the late-glacial) to less extreme winter-biased groundwater recharge
ratios (e.g., forests during late-Holocene; Jasechko et al., 2014).
South America
Measured Δ18Olate-glacial values across South America
range from -6.3 to +0.6 ‰
(Figs. 2 and 6). The highest-magnitude, negative measured Δ18Olate-glacial values are found in Andean ice cores (Δ18Olate-glacial of -4.6 ± 1.0 and -6.3 ± 1.3;
Thompson et al., 1995, 1998). Here the importance of upstream convection
upon modern Andean precipitation δ18O has been highlighted at
inter-annual (Hoffmann et al., 2003; Vuille and Werner, 2005), seasonal
(Vimeux et al., 2005; Samuels-Crow et al., 2014) and daily timescales
(Vimeux et al., 2011). It is therefore possible that upstream convection
controls past changes to Andean precipitation isotope compositions recorded
in ice cores.
The measured groundwater Δ18Olate-glacial value located
in eastern Brazil is -2.7 ± 1.3 ‰ (Salati et al.,
1974). Eastern Brazil was 5 ∘C cooler than today during the latter
half of the last glacial period (Stute et al., 1995b). Four of the five
general circulation models simulate positive Δ18Olate-glacial values across eastern Brazil (Fig. 5),
highlighting a difference between simulated and measured Δ18Olate-glacial values in parts of the tropics. The negative
measured Δ18Olate-glacial value in eastern Brazil has
been previously interpreted to reflect higher-than-modern precipitation
during the last glacial time period (Salati et al., 1974). Lewis et al. (2010) show that localized rainfall governs precipitation δ18O
in eastern Brazil. Modern precipitation δ18O values are lowest
in eastern Brazil when precipitation rates are at a maximum. Extending Lewis
et al.'s interpretation linking local precipitation amount to precipitation
δ18O would suggest that the negative measured Δ18Olate-glacial value found in eastern Brazil may indeed
record wetter-than-modern conditions during the late-glacial as proposed
by Salati et al. (1974). Further, disagreement between measured and
simulated Δ18Olate-glacial in eastern Brazil highlights
the need to critically evaluate climate model performance in regions where
the precipitation amount is closely correlated with precipitation δ18O.
North America
Measured Δ18Olate-glacial from North American proxy
records range from -5.5 to
+1.0 ‰. Canadian records of groundwater recharge that
took place beneath the Laurentide ice sheet are not included in this
synthesis (“subglacial recharge”; Grasby and Chen, 2005; Ferguson et al.,
2007; McIntosh et al., 2012; Ferguson and Jasechko, 2015). These records
were excluded because the subglacial meltwaters that recharged aquifers
likely reflect precipitation that fell elsewhere on the paleo-ice sheet,
potentially complicating the comparison of groundwater isotope compositions
for the late-Holocene and last glacial time period.
Measured Δ18Olate-glacial values along the USA east
coast show the highest, positive values in Georgia (latitude: 32∘ N; measured Δ18Olate-glacial of
+1.0 ‰; Clark et al., 1997), decreasing northward to
near-zero measured Δ18Olate-glacial values in coastal
Maryland (latitude 39∘ N; measured Δ18Olate-glacial of -0.1 ± 0.4 ‰;
Aeschbach-Hertig et al., 2002). Decreasing Δ18Olate-glacial values with increasing latitude along the USA
east coast may be explained in part by the isotopic distillation of air
masses advected northward from the subtropics under cooler-than-modern final
atmospheric condensation temperatures. Indeed, paleoclimate records indicate
that Maryland was more arid and as much as 9–12 ∘C cooler during
the late-glacial relative to the late-Holocene (Purdy et al., 1996;
Aeschbach-Hertig et al., 2002; Plummer et al., 2012). In addition to
temperature change, late-glacial precipitation isotope compositions along
eastern USA coastline were likely impacted by the lower-than-modern
late-glacial sea levels, which changed overland atmospheric transport
distances between the late-glacial and late-Holocene (Clark et al., 1997;
Aeschbach-Hertig et al., 2002; Tharammal et al., 2013).
Measured Δ18Olate-glacial values in the central and
southwestern USA have the highest magnitude, negative measured Δ18Olate-glacial values of temperate North America, ranging
from -1.0 to -3.4 ‰ . Central
and southwestern USA measured Δ18Olate-glacial values
contrast the positive measured Δ18Olate-glacial values
found along the eastern USA coast at similar latitudes. Consistently
negative Δ18Olate-glacial values in central and
southwest USA suggest that advected moisture to the region underwent greater
upstream air mass distillation during the late-glacial than under modern
climate. Pollen, vadose zone and groundwater records show that
late-glacial southwestern USA was ∼ 4 ∘C cooler,
had greater groundwater recharge fluxes, and had more widespread forests
than present day (Stute et al., 1992, 1995a; Scanlon et al., 2003; Williams,
2003). Negative measured Δ18Olate-glacial values found
in the southwest USA have been ascribed to lower-than-modern summer
precipitation (New Mexico, Phillips et al., 1986), latitudinal shifts in the
positions of the polar jet stream and the intertropical convergence zone
(New Mexico, Asmerom et al., 2010) and changes to over-ocean humidity,
temperature or moisture sources (Idaho, Schlegel et al., 2009). Wagner et al. (2010) interpret decreases to southwestern precipitation δ18O to reflect cooler and more-humid conditions. Extending this
interpretation to negative measured Δ18Olate-glacial
values found across the southwestern USA values supports earlier conclusions
that the region was cooler and more humid than today during the
late-glacial, possibly linked to changes in air mass trajectories and
moisture sources (Asmerom et al., 2010; Wagner et al., 2010). Simulated
Δ18Olate-glacial values across North America closely
match spatial patterns of measured Δ18Olate-glacial
synthesized in this study. Strong, multi-model agreement with measured
Δ18Olate-glacial patterns supports continued
application of isotope-enabled general circulation models when interpreting
North American precipitation isotope proxy records.