Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30
https://doi.org/10.1186/s40645-022-00489-0

Progress in Earth andPlanetary Science

Open Access

REVIEW

The response of the hydrological cycle
to temperature changes in recent and distantclimatic history
Shailendra Pratap*   and Yannis Markonis

Abstract
The relationship between the hydrological cycle and the temperature is rather complex and of great importance tohuman socioeconomic activities. The prevailing theory suggests that as temperature increases the hydrological cycle
is intensified. Practically, this means more and heavier precipitation. However, the exact magnitude of hydrologicalcycle response and its spatio-temporal characteristics is still under investigation. Looking back in Earth’s hydroclimatic
history, it is easy to find some periods where global temperature was substantially different than present. Here, we
examine some of these periods to present the current knowledge about past hydrological cycle variability (specifically
precipitation), and its relationship to temperature. The periods under investigation are the Mid-Miocene Climate Optimum, the Eemian Interglacial Stage, the Last Glacial Maximum, the Heinrich and Dansgaard–Oeschger Events, theBølling–Allerød, the Younger Dryas, the 8.2 ka event, the Medieval Climate Anomaly, and the Little Ice Age. We reportthat the hypothesis that a warmer climate is a wetter climate could be an oversimplification, because the response ofwater cycle appears to be spatio-temporally heterogeneous.
Keywords: Global water cycle, Paleoclimate, Hydrological cycle, Water cycle intensification, Hydroclimatic variability
1 Introduction
Looking back in Earth’s hydroclimatic history, there havebeen substantial shifts in the hydrological cycle (Ljungqvist et al. 2016). In the past few million years, many
rapid climate transitions have occurred, with time scales
ranging from decades to centuries (Corrick et al. 2020).
For example, during Holocene, i.e. the last 18–20 thousand years (ka) before present (BP), the paleoclimaticrecords show considerable fluctuations in both the seasonal and spatial distribution of precipitation (Badgeley
et al. 2020). During the late glacial (18–16.5 ka), sea surface temperature (SST) was about 5–10 °C colder than
the recent Holocene (11.5–9 ka) over both the North
Pacific and the North Atlantic (Praetorius et al. 2020).
For the same period, the global averaged precipitation*Correspondence: pratap@fzp.czu.cz
Faculty of Environmental Sciences, Czech University of Life Sciences Prague,Kamýcká 129, Praha – Suchdol 165 00, Czech Republic

was about 10–14% lower than today, with the maximum
reduction over the Northern Hemisphere (NH) due toreduced convective activity (Gates 1976; Kwiecien et al.
2009; Sun et al. 2019). As the Last Glacial ended andthe climate became warmer, there was a shift to wetter
conditions as well. From 13 to 12 ka BP, the monsoon
circulation was intensified, resulting to an increase inprecipitation by about 200–300 mm at lower latitudes(Knox and Wright 1983; Maher 2008; Pausata et al. 2020).
Stronger monsoons were also observed between 8 to 3 ka
BP, coupling the widespread warming (Chawchai et al.
2021). The most affected region was East Asia (Rao et al.
2016), where precipitation was over 30% higher than
today from 7.8 to 5.3 ka BP (Yang et al. 2016). All thesechanges occurred in various spatiotemporal scales, andtherefore, it is still challenging to estimate the hydrological cycle variability and quantify it on global, continental,
and regional scales.

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Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Besides natural variability, anthropogenic forcing(GHG emissions and land-use changes) is also regarded
as one of the possible drivers of abrupt changes of thehydrological cycle (Allan et al. 2014). Global warming
is expected to intensify the global hydrologic cycle, andincrease the frequencies of extremes like heavy rainfall,
flood, and drought conditions (Huntington 2010). The
term intensification of the hydrological cycle is used todescribe an acceleration in the rates of atmospheric water
vapor content, precipitation, evaporation, and evapotranspiration (ET) (Trenberth 1998). There is a solid theoretical basis that relates atmospheric warming and theintensification of hydrological cycle. The basis for thisrelationship is the Clausius–Clapeyron relation, whichsuggests the exponential response of specific humidity
to temperature increase (Huntington 2006). The Clausius–Clapeyron formulation implies the slope of thisrelationship has to remain below the 6.5% per Kelvin as
the evaporation is energy limited (Norris et al. 2019).
However, precipitation is also energy limited, because theatmosphere should be able to radiate away the latent heatproduced during precipitation events (O’Gorman 2012).
This makes the estimation of the precipitation response
under energy constrain conditions a complex task.
Due to this complexity, General Circulation Models(GCMs) are being extensively used in the estimation of
the intensification hydrological cycle (Watterson et al.
1997). The GCMs still show strong variance in theirresults, although there is general agreement that there isa detectable increase in global mean precipitation, also
evident in observational records (Markonis et al. 2019).
For example, Allen and Ingram (2002) reported that the
precipitation will increase by approximately 3.4% per
Kelvin degree, while Wentz et al. (2007) report a slower
rate, between 1 and 3% per Kelvin. Another study using
20 coupled ocean-land–atmosphere models showsthat precipitation, runoff, and evaporation will globally
increase by 5.2%, 7.3%, and 5.2%, respectively, responding to a mean surface air temperature increase of 2.3 °C
by 2050 (Wetherald and Manabe 2002). Durack et al.
(2012) present a 4% increase in precipitation in response
to 0.5 °C warming. As we see, the precise quantificationof the relationship between temperature and hydrological cycle remains unresolved. A plausible alternative andcomplementary approach to the GCMs is the study of thepast states of hydrological cycle through paleoclimaticreconstructions. By investigating the past hydroclimatic
variability range, we can shed more light to the underlying physical mechanisms and/or constrain the climate
model simulations (Seftigen et al. 2017).
This study aims to map the current knowledge abouthydrological cycle variability, and its relationship to temperature. Since it is extremely difficult to assess all the

Page 2 of 37

processes related to the global hydrological cycle, we
focus our review to precipitation and temperature (asaproxy for evaporation), which can be used indirectly
to describe the global water balance Vargas Godoy et al.
(2021). We have selected past periods with significant
hydroclimatic fluctuations, that span from centuries tomillion years. The lengthiest of them is Mid-Miocene
Climate Optimum (MMCO; 17–14.5 million years BP),when global temperature was 3–8 °C higher than preindustrial levels. Such a warmer climate can help us
determine future changes of water cycle to extremely
high temperature conditions. Alternatively, warmer
periods such as the Eemian Interglacial Stage (temperature 1.3 °C higher than today) can provide insight in thenear future changes due to global warming. On the other
hand, the study of ice age climates can help us determine
the hydrological cycle response to colder regimes (e.g.
Last Glacial Maximum when global temperature was
4.3 °C lower than today). The rapid transitions between
cold and warm conditions are also of interest, and here,
we will explore the hydrological cycle shifts during theHeinrich and Dansgaard–Oeschger Events. Finally, thestudy of Holocene allows us to examine time scales closer
to the one of the recent temperature increase. We investigate the hydroclimatic conditions for Bølling–Allerød,
Younger Dryas, the 8.2 ka event, the Medieval Climate
Anomaly, and the Little Ice Age. Assessing the state ofhydrological cycle during all these periods can offer analternative pathway for anticipating the hydroclimaticchanges that are yet to come both in the near and distantfuture (Meehl et al. 2007).
Please note that the pre-industrial values of temperature or precipitation corresponds to the period 1850–
1900. On the other hand, there are studies that compare
the climatic conditions with today. In this case, we
assume today as the reference time when the corresponding study was published (industrial era). We use the same
assumption for the studies without any explicit reference
to a comparison period.

2 Climatic regimes of the distant past
2.1 MMCO

The MMCO (14 million years BP) is a rather long periodof significantly warmer conditions compared to present(Böhme et al. 2007). What makes it particularly interesting is the evidence of enhanced fluctuations in the carbon
cycle (Holbourn et al. 2014). Proxy records of alkenones(Zhang et al. 2013), paleosols (Breecker and Retallack2014), stomata (Grein et al. 2013), and marine boron isotopes (Greenop et al. 2014) show that during the MMCO
event, atmospheric CO2 was less than 450 ppm, which is
not far from the current CO2 levels and within the rangeof near future CO2 projections (Steinthorsdottir et al.

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

2020). However, there are also studies that report lower
CO2 concentrations, equal to or less than today (Zhang
et al. 2013), implying that CO2 might not be the main climatic driver (Pearson and Palmer 2000). Nevertheless,
MMCO presents an excellent opportunity to investigate
the functioning of the hydrological cycle in a warmer
climate.
Air temperature reconstructions and model simulations suggest that during the MMCO, the annual mean
global temperature was between 3 and 8 °C more thanthe pre-industrial levels (You et al. 2009; Pound et al.
2012; Steinthorsdottir et al. 2020). This is in good agreement with temperature proxies of deep-ocean water,which reveal a 5–6 °C warmer temperature as of today(Haq 1973; Miller et al. 1991; Zachos et al. 2008). The
regions with higher temperatures are located mostly
at mid to high latitudes (Böhme et al. 2007; Bruch et al.
2007; You 2010). Alongside with the warmer conditions, the MMCO also exhibited a rather humid climate(Zachos et al. 2001). This is also supported by model
simulations, which show widespread increases in meanannual precipitation across northern and central Africa,
North America, northern Eurasia, and Greenland (Kennett 1994; Fox and Koch 2004; Retallack 2007; Henrot
et al. 2010; Herold et al. 2011).
The prevailing wet conditions are also confirmed byregional studies. Wet conditions of the MMCO have also
been reported for Europe, where there was an increasein average annual precipitation of about 830–1350 mm(Böhme et al. 2007; Methner et al. 2020; Kuhlemann
and Kempf 2002; Schlunegger et al. 1996). In addition,
the isotope estimations at the Pannonian basin (Central
Europe) suggest higher summer precipitation during theLate Miocene (about 10 million years BP) (Harzhauser
et al. 2007). Pollen and leaf proxies from the Nenanacoal
field (Alaska Range, Alaska) imply a particular warmer
period from about 18 to 14 million years BP (Leopold
and Denton 1987). Pollen investigation at the Tian Shan
(China) and sediment analysis at northeastward ofTibetan Plateau (China) show a wet and warm stage (Sunand Zhang 2008; Song et al. 2018). Stable isotope sclerochronology over northern South America (Guajira Peninsula, Colombia) indicates wet conditions with enhanced
seasonality in regions that today have semiarid conditions due to a northerly shift of the Inter Tropical Convergence Zone (ITCZ) (Scholz et al. 2020). Finally, warm
and wet climate dominated at Antarctica and the someregions of Southern Hemisphere (SH) high latitudes (You2010; Feakins et al. 2012).
We have to note, though, that there is also evidence forincreased aridity over Africa (Retallack 1992; Levin et al.
2006; Eronen et al. 2012; Morales-García et al. 2020),
Australia (Stein and Robert 1985; Byrne et al. 2008; Wu

Page 3 of 37

et al. 2018), South America (Pascual and Jaureguizar
1990), and some regions of North America (Wolfe
1985; Chamberlain et al. 2014) and Asia (Jiang and Ding
2010; Liu et al. 2009). In the latter, there was an expansion of the arid region from the western to the eastern
coast of China, whereas the humid areas were limited to
the northern and southern parts (Steininger 1999; Wanet al. 2007; Clift et al. 2014). The physical mechanism that
regulated the aridification over Asia, and the widespread
mid-latitude arid region of the NH remains enigmatic(Hou et al. 2014). Analysis of bulk δ13C, over the centraleastern Idaho (Railroad Canyon section, USA), suggests
an average mean annual precipitation of about 190 mm(ranges from 10 to 510 mm/year) during the MMCO thatis almost equivalent to today’s values (about 236 mm/year) (Harris et al. 2020). In addition, a paleosols analysis
over the northern Pakistan (Zaleha 1997) suggests middle Miocene monsoon was similar to today (Allen andArmstrong 2012).
In Table 1, all the analysed studies are presented byregion, hemisphere, latitudinal zone, and time period. Inorder to highlight the spatiotemporal variability of thehydroclimatic conditions, some studies appear to morethan one rows, e.g. You (2010). In this manner, we can seethat even in a much warmer world, there is no uniform
shift of hydrological cycle; some regions became wetter, some became drier, and some appear to be similar to
today. Still, the comparative examination of temperatureand precipitation reveals that warmer conditions favor
more an increase in precipitation than drier climate conditions in an approximately 2:1 ratio (Fig. 1).
2.2 The Eemian Interglacial StageThe Eemian Interglacial Stage, also known as the Marine
Isotope Stage (MIS) 5e, is a period that lasted 15 to 17
thousand years at approximately 130 ka BP (Abarbanel
and Lall 1996). Commonly referred as the Last Interglacial, it is the period that preceded the last glacial stage,
with stable climatic conditions similarly to Holocene. Initially, the Eemian was thought to be quite warmer thaninterglacial. Andersen et al. (2004) reported that thetemperature was 5 °C higher as to today, according to anoxygen isotope reconstruction. However, more recent
studies suggested that global average surface temperature was up to 1.3 °C warmer than the pre-industrial levels (Fischer et al. 2018), reaching a 2 °C maximum in the
middle of the period (Snyder 2016). The global averagetemperature over land was 1.7 °C warmer than the preindustrial levels, while the oceans were 0.8 °C warmer(Otto-Bliesner et al. 2013). The temperature differenceswere quite heterogeneous over land. The mid and highnorthern latitudes experienced considerably warmer
temperatures, ranging between 2 and 5 °C (Turney andPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 4 of 37

Table 1 Temperature and precipitation conditions during the MMCO
Study site

Hemisphere

Zone

PeriodT

P

Citation(s)GlobalNH/SH

0

17–14.5Warm

0You et al. (2009)
Pound et al. (2012)
Steinthorsdottir et al. (2020)

N. Hemisphere

NH

M-L

17–14.5

Warm

Dry

Böhme et al. (2007);
Bruch et al. (2007); You (2010) Zachos et al. (2001)

N. Hemisphere

NH

H-L

17–14.5

Warm

Dry

Böhme et al. (2007)
Bruch et al. (2007), You (2010), Zachos et al. (2001)

Deep Ocean

NH

H-L

17–14.5Warm0

Haq (1973), Miller et al. (1991)
Zachos et al. (2008)

S. Hemisphere

SH

H-L

17–14.5

Warm

Wet

You (2010), Feakins et al. (2012)

Asia

NH

M-L

17–14.50

Dry

Jiang and Ding (2010)
Liu et al. (2009)

AfricaNH

L-L

17–14.50

DryRetallack (1992)
Levin et al. (2006)
Eronen et al. (2012)
Morales-García et al. (2020)

Antarctica

SH

H-L

17–14.5

Warm

Wet

You (2010); Feakins et al. (2012)

S. America

SH

L-L

17–14.5

0

Dry

Pascual and Jaureguizar (1990)

N. America

NH

M-L

20–14.5

Warm

Wet

Kennett (1994); Retallack (2007)
Fox and Koch (2004); Henrot et al. (2010) Herold et al. (2011)

N. America

NH

M-L

17–14.5

0

Dry

Wolfe (1985) Chamberlain et al. (2014)

EuropeNH

M-L

17–14.5

0Wet

Böhme et al. (2007); Methner et al. (2020)
Kuhlemann and Kempf (2002)
Schlunegger et al. (1996)

AustraliaSH

M-L

17–14.5

0

Dry

Stein and Robert (1985)
Byrne et al. (2008)
Wu et al. (2018)

N. Eurasia

NH

H-L

21–14.5

Warm

Wet

Kennett (1994); Retallack (2007)
Fox and Koch (2004); Henrot et al. (2010) Herold et al. (2011)

Central Africa

NH

L-L

19–14.5

Warm

Wet

Kennett (1994); Retallack (2007)
Fox and Koch (2004); Henrot et al. (2010) Herold et al. (2011)

N. AfricaNH

M-L

18–14.5

Warm

Wet

Kennett (1994); Retallack (2007)
Fox and Koch (2004); Henrot et al. (2010) Herold et al. (2011)

Northern S. America

SH

L-L

17–14.50

Wet

Scholz et al. (2020)

China

NH

M-L

17–14.5

Warm

Wet

Sun and Zhang (2008)
Song et al. (2018)
Steininger (1999)
Wan et al. (2007)
Clift et al. (2014)

GreenlandNH

H-L

21–14.5

Warm

Wet

Kennett (1994); Retallack (2007)
Fox and Koch (2004); Henrot et al. (2010) Herold et al. (2011)

E. Idaho, USA

NH

M-L

17–14.50

DryHarris et al. (2020)

N. Pakistan

NH

M-L

17–14.50

Dry

Allen and Armstrong (2012)

Alaska Range, Alaska

NH

H-L

18–14

Warm

0

Leopold and Denton (1987)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes. Period units are in million years BP and average age
uncertainty is ± 1 million years

Jones 2010), which are comparable to some global warming projections (Change 2014). Similarly to the MMCO,
the Eemian is also an excellent analogue for analysing thestate of hydrological cycle in warmer conditions (Adams
et al. 1999).
Most of the available paleoclimatic records show that
the Last Interglacial was wetter than Holocene. This isalso supported by model simulations, demonstrating an

intensified hydrological cycle (Weaver and Hughes 1994;
Pedersen et al. 2017; Johnston et al. 2018; Zhang et al.
2021b). Enhanced precipitation is observed mainly at the
NH in paleoclimatic records over the low latitudes (Williams et al. 2020), boreal mid-latitude regions (Members
2006), and the Arctic (Kim et al. 2010). In addition, theice melt pulses from Greenland have been suggested toinfluence the enhanced climate variability across thePratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 5 of 37

Fig. 1 Relationship of temperature and precipitation during the MMCO. The number of studies used for the warm/cold or wet/dry conditions canbe found in Table 1

Mediterranean (Tzedakis et al. 2018). During that time,
when insolation was at its peak over the NH (Nehme
et al. 2015), wet intervals were observed over Southern
Europe (Brauer et al. 2007), and specifically, over the
Eastern Mediterranean (Bar-Matthews 2014; Bar-Matthews et al. 2019). Continental North America was also
wetter and warmer compared to today (Anderson et al.
2014). However, there were, also, some fluctuations to
dry intervals (Curry and Baker 2000), which are further
observed in the high values of carbon isotope (δ13C29 andδ13C31) (Suh et al. 2020).
Furthermore, there is an increase in NH summer monsoons (Wang et al. 2008). Both the proxy and model
approaches explicitly suggest higher monsoon activity
over North African and Asia (Prell and Kutzbach 1987;
Scussolini et al. 2019). Terrestrial proxy records suggest wetter and warmer climate over the Sahara Arabiandesert area compared to the present (Rosenberg et al.
2013; Petit-Maire et al. 2010). This is further confirmed
by both the oxygen isotopes on speleothems at Soreq
Cave (Israel) and climate models, showing increased
regional rainfall during the Last Interglacial, attributed towetter winters and increased summer monsoons (Orland
et al. 2019). In addition, speleothems and fossil corals

reconstructions in the reef terraces also indicate a wetter
Eemian interglacial alongside the Gulf of Aqaba at Arabian Peninsula (Yehudai et al. 2017). Similar speleothem
findings as well confirm a wetter climate over Southern
Arabia (Vaks et al. 2006).
On the other hand, there are also regions that experienced enhanced aridity. The evaluation of the Eemian climate across Europe using pollen reconstructions presents
a different picture to the one described above. Colder anddryer conditions prevailed in the southern regions andconditions that are similar to today in the higher latitudes (Brewer et al. 2008). Sediment records from Maarlake (Germany) show a late Eemian cold and arid eventthat lasted 468 years (Sirocko et al. 2005). Weakening of
the southern summer monsoon has been reported in themodeling and some proxy records (Montoya et al. 2000).
Supporting evidence can be found in the speleothems ofWestern Australia, which indicate arid conditions (Zhaoet al. 2001). Drier conditions also appeared in Argentina as detected in loess (paleosols) records (Tofalo et al.
2011), and Bolivia, where sediment records from Lake
Titicaca suggest warmer and more arid conditions during the Eemian period (Fritz et al. 2007). This seems to
be a recurring pattern during warm interstadials andPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

interglacials, when the southeastern regions of Australia
show comparatively arid conditions (Ayliffe et al. 1998).
In general, both proxy records and model simulations
suggest weakened monsoonal precipitation over the SH
compared to the pre-industrial times (Nikolova et al.
2013).
Similarly to the MMCO, during the MIS-5e, there wasa substantial warm-and-wet pattern which was far fromhomogeneous. The majority of the studies on temperature shows warmer climate, i.e. about 85%, while the rest
reveal cold conditions (Fig. 2 and Table 2). In precipitation records, the difference is slightly milder with about
75% of the studies suggesting wet conditions and about
25% a drier climate.
2.3 The Last Glacial Maximum

The Last Glacial Maximum (LGM) corresponds to theperiod during the last Glacial Stage that the ice sheets
extended to their maximum length reaching their highest mass. It occurred between 30 and 15 ka (Prentice
et al. 1992), although more recent estimates place it
between 26.5 and 19 ka BP (Clark et al. 2009). During
the LGM, the climate conditions at NH high latitudes
were much colder and drier than today (Bigelow et al.

Page 6 of 37

2003; Otto-Bliesner et al. 2006; Yokoyama et al. 2000).
The global average temperature is estimated at 3–6 °C
lower than the modern values (Bush and Philander
1999; Schmittner et al. 2011), while locally, e.g. at Greenland Summit, reached approximately 15–20 °C colder
than the present levels (Johnsen et al. 1995; Cuffey et al.
1995; Miller et al. 2010). Even the tropics were substantially colder, ranging between 2 and 3.5 °C below present
temperatures (Barker et al. 2005; Annan and Hargreaves2013). This has also been confirmed by model results,
which also estimate the difference around 2.5 °C acrossthe equatorial regions (Crowley 2000; Ballantyne et al.
2005). Similarly to the Eemian the main driver for thetemperature decline is the incoming insolation (Bush andPhilander 1999; Clark et al. 2009).
The decline in temperature is also confirmed bydecrease in the SST over multiple oceans. The Multiproxy Approach for the Reconstruction of the Glacial
Ocean surface (MARGO) project suggests that there
was an annual tropical SST cooling of 1.7(± 1) °C during
the LGM. Similarly, the eastern and western equatorial
Pacific, northwestern Pacific subarctic gyre, and northwestern tropical Pacific regions also show that the SST
was lower (0.9–3.6 °C) than the present (Kucera et al.

Fig. 2 Relationship of temperature and precipitation during the Eemian Interglacial Stage. The number of studies used for the warm/cold or wet/dry conditions can be found in Table 2

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 7 of 37

Table 2 Temperature and precipitation conditions during the Eemian Interglacial Stage
Study site

Hemisphere

Zone

PeriodT

P

Citation(s)GlobalNH/SH

0

130–116Warm

0Fischer et al. (2018)
Snyder (2016)
Otto-Bliesner et al. (2013)

GlobalNH/SH

0

130–1160

Wet

Weaver and Hughes (1994)
Pedersen et al. (2017)
Johnston et al. (2018)
Zhang et al. (2021b)

N. HemisphereNH0

130–116Warm

0Andersen et al. (2004)

N. HemisphereNH0

130–1160

Wet

Wang et al. (2008)
Williams et al. (2020)
Members (2006)
Kim et al. (2010)

N. Hemisphere

NH

M-L

130–116Warm

0Turney and Jones (2010)

N. Hemisphere

NH

H-L

130–116Warm

0Turney and Jones (2010)

Asia

NH

M-L

130–1160

Wet

Prell and Kutzbach (1987)
Scussolini et al. (2019)

N. America

NH

M-L

130–116

Warm

Wet

Anderson et al. (2014)

EuropeNH

M-L

130–116

Cold

Dry

Brewer et al. (2008)

Australia

SH

M-L

130–116

0

Dry

Ayliffe et al. (1998)

N. AfricaNH

M-L

130–1160

Wet

Prell and Kutzbach (1987)
Scussolini et al. (2019)

S. EuropeNH

M-L

130–1160

Wet

Brauer et al. (2007)

W. Australia

SH

M-L

130–1160Dry

Zhao et al. (2001)

Arabian desert

NH

M-L

130–116

Warm

Wet

Rosenberg et al. 2013 Petit-Maire et al. (2010)

GreenlandNH

H-L

130–116Warm

0Tzedakis et al. (2018)

E. Mediterranean

NH

M-L

130–1160

Wet

Bar-Matthews (2014) Bar-Matthews et al. (2019)

Arabian Pen

NH

M-L

130–1160Wet

Yehudai et al. (2017)

GermanyNH

M-L

130–116

Cold

Dry

Sirocko et al. (2005)

BoliviaSH

L-L

130–116Warm

Dry

Fritz et al. (2007)

Argentina

SH

M-L

130–116

0

Dry

Tofalo et al. (2011)

Argentina

SH

M-L

130–116

0

DryNikolova et al. (2013)

S. Arabian Pen

NH

M-L

130–116

0

Wet

Vaks et al. (2006)

Soreq Cave, Israel

NH

M-L

130–1160

Wet

Orland et al. (2019)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes. Period units are in ka BP and average age uncertainty
is ± 5 ka

2005). The lower SST resulted in increased upwelling ofcolder water across the continental margin and, finally,a cooler climate, especially over the NH (Rosell-Mel´eet al. 2004). There is also limited evidence about the
SST decline in finer scales. For example, the Mediterranean Sea shows that the SST was about a 1 °C lower
than the present, particularly in the eastern part (Hayes
et al. 2005). On the other hand, not all the studies agree
on a lower SST during the LGM. The SST derived fromthe central tropical Pacific and northern subtropics weresimilar to the modern levels of the SST (Lee et al. 2001),
while a few regions have experienced a higher SST, suchas the Northwest Pacific margin, southern parts of Iceland–Faroe Ridge, Iberian margin, north–south-west
African boundary currents, and Japan Sea (Waelbroeck

et al. 2009). Still the majority of the SST records advocate for cold conditions, which are expected to affectthe hydroclimate of the nearby landmasses (Seager et al.
2007).
Most of the proxy records suggest that during the LGM
the global hydrological cycle was weaker compared to
today (Cragin et al. 1977; Yung et al. 1996; Steffensen
1997; Li and Zhang 2020). Dry conditions were typicalover both hemispheres and model simulations show thatthe decline in global temperature is linked to a decline inatmospheric water vapor concentration. Otto-Bliesner
et al. (2006) estimated that precipitable water was 18%
less than today resulting to an annual average precipitation of about 2.49 mm per day. The weakening the globalhydrological cycle is due to a reduction of about 10% inPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

both evaporation and precipitation (Bush and Philander
1998; Gasse 2000). The simulations also suggested a surplus of precipitation over evaporation that has lowered
the net amount of water vapor in the atmosphere (Bushand Philander 1999; Rojas et al. 2009).
Proxy records and model simulations (CCSM3) reporta weakened summer monsoon for both tropical as well
as northern Africa (Prentice et al. 2000). Moreover, analysis of lake sediments from the Pretoria Saltpan (South
Africa) suggests a negative shift in the monsoonal precipitation with total precipitation approximately 15 to
20% less than today (Patridge et al. 1997; Simon et al.
2015). The drier conditions were also confirmed by diatom estimates from the same site (Metcalfe 1999; Gasseand Van Campo 2001). The lake records from the eastand southwest Amazonia also suggest lower precipitationlevels than the present (Absy et al. 1991; Sifeddine et al.
2001). Similar changes are reported for high latitudes.
Lake sediment records over southern east Siberia (Lake
Baikal) show a drop of about 11% in annual precipitation
and about 80% drop in summer precipitation, comparedto the present climate (Osipov and Khlystov 2010). Similarly, the yearly precipitation over the Greenland Summit
has been found up to three times less than the present
values (Cuffey and Clow 1997; Johnsen et al. 2001).
In Europe, where regional climate modeling suggests that the annual average air temperature was about
6–9 °C lower than the present, while the precipitation
was quite lower, especially over the northern regions(Strandberg et al. 2011). Interestingly, the decline waslinked to a change in the atmospheric circulation pattern that determines the precipitation regime andstrength. Currently, the precipitation pattern over central Europe is controlled by a westerly to northwesterly
circulation system. During the LGM, the atmosphericmoisture reached central Europe through south-westerly advection (Becker et al. 2016). This was also supported by the oxygen isotope analysis on speleothemsof the Sieben Hengste cave (Bernese Alps), which reportsouthwesterly moisture advection (during 26.5–23.5 ka)(Luetscher et al. 2015). The change in atmospheric circulation resulted to an increase in precipitation over southern Europe (Kuhlemann et al. 2008). In the eastern andcentral Mediterranean, there has been evidence for an
increase in mountain glaciers at several locations, as wellas an increased rate of winter precipitation (Strandberg
et al. 2011). The Mediterranean is not the only region
that wetter conditions appeared, as there is evidence ofsimilar fluctuations over the extra-tropics (Clark andMix 2002). Similar findings were found in the assessment
on lake levels over East Africa, even though palaeovegetation analysis point to a dry climate (Barker and Gasse
2003). However, these changes are spatially limited andPage 8 of 37

do not significantly alter the global signal of decline inprecipitation.
Contrary to the warm conditions of the MMCO andthe Eemian Interglacial Stage, the cold conditions thatprevailed in the LGM are mostly associated with drier
climate. However, again the climatic conditions may differ spatially. For instance, in Fig. 3 and Table 3, we cansee that about 20% of precipitation records correspondto regions with wet climate during this doubtlessly cold
period. Another plausible explanation, besides spatiallyheterogeneity, could lie to the climatic proxy nature andthe processes involved, which might falsely interpretsolid precipitation or glacial extension as a wet regime.
In any case, the global signal advocates for a weakening inwater cycle strength (Li and Zhang 2020).

3 The abrupt climatic events of the last glacial
3.1 Dansgaard–Oeschger and Heinrich events

During the last glacial period, Earth’s climate has gone
through some abrupt changes over the North Atlantic region (Dansgaard et al. 1993). Proxy records suggest more than 24 cooling and warming events, termed
as the Dansgaard–Oeschger (D–O) events (Rasmussen
et al. 2016). During the D–O events, most of the NH is
influenced by abrupt warming, which is then succeededby a more gradual cooling (Martrat et al. 2004). Ice
core records collected from Greenland suggest a rapid
increase in atmospheric temperature ranging between
10 and 16 °C that occurred within a few decades (Johnsen et al. 1989; Lang et al. 1999; Budsky et al. 2019). In
addition, there is evidence that warmer climate conditions were coupled with higher precipitation (Genty et al.
2003). The factors driving the D–O events are undervigorous debate, ranging from ocean–atmosphere orsea ice-atmosphere interactions (Broecker et al. 1990;
Li and Born 2019) to cyclic Greenland ice sheet calving(Van Kreveld et al. 2000) and Earth’s orbital forcing (Van
Geel et al. 1999). Widespread signs of D–O events in theNordic seas and North Atlantic have been found to be
associated with the Atlantic Meridional Overturning Circulation (AMOC) instability, influenced by the variability in convection rate (Rasmussen et al. 2016). However,there are also D–O events that did not only influence
the North Atlantic, but had a large-scale, or even global,
impact to the climatic system. The fingerprint of D–O
events can be found in deep-sea records, where it can beseen on planktonic and benthic records across the globe(Shackleton et al. 2000), or the Vostok ice core record at
Antarctica (Jouzel et al. 1987). This is probably due to therelationship between the D–O events and the intensity of
the AMOC (Santos et al. 2020).
The D–O events are also evident over the Mediterranean region, where there was an increase in precipitation

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 9 of 37

Fig. 3 Relationship of temperature and precipitation during the LGM. The number of studies used for the warm/cold or wet/dry conditions can befound in Table 3

over the Iberian Peninsula (Nebout et al. 2002; Budsky
et al. 2019) and Italy (Alley and Clark 1999). In addition, the D–O events are observed in the oxygen isotope
record of the Soreq cave (Israel), where low δ18O andhigh δ13C values suggest wet conditions (Bar-Matthews
et al. 2000). An increase in precipitation is also reported
for Great Basin (Nevada; United States), where there is anincrease in the lake levels, derive by the analysis of δ18O
proxy records (Benson et al. 1998). Some D–O events canalso be linked with climatic fluctuations across the Indian
Ocean (Altabet et al. 2002), such as events D–O events
7 and 8, which occurred at approximately 34–41 ka BP(Beck et al. 2001). The δ18O estimates in the stalagmites
collected from northern Vietnam, Indian, and Chinese
caves show strengthening in the Indian and Asian summer monsoons (Dung et al. 2020; Cheng et al. 2016;
Kathayat et al. 2016). Additionally, the D–O event 12(45 ka BP (Genty et al. 2003)) was linked to the increased
intensity of the Asian southwest monsoon during about
50–40 ka (Anderson and Prell 1993). Similar findings
have been reported for other regions over Asia (Wanget al. 2001), while there is evidence that the D–O events
can also be detected at South America (Peterson et al.
2000).

Between the D–O events, there are also some abrupt
transitions to rather cold periods. They were named afterHartmut Heinrich, who investigated the characteristics of
six intervals from 70 to 14 ka BP that occurred between
the D–O events and appear to be the coldest events ofthe glacial (Heinrich 1988). The Heinrich events affectedmost of the Eurasia and North America, resulting todrier and colder conditions (Genty et al. 2003; Benson
et al. 1996; Asmerom et al. 2010). Although the driversof the Heinrich events are still not fully understood, there
is general agreement that they are related to changes in
the oceanic circulation over the North Atlantic (Thomas
et al. 1995) and in the ice sheets over NH (Broecker
2000). They are mainly linked with the release of large
volume of freshwater through iceberg melting (Boers
2018). These large-scale cold freshwater pulses caused
further changes all over the global climatic system.
A 5 to 8 °C cooling has been observed over the Mediterranean surface water (Rohling et al. 1998), and significant aridity has been observed over the southwestern
USA (Wagner et al. 2010). The influence of some Heinrich events extends to the tropics, where enhanced aridity has been reported (Leuschner and Sirocko 2000).
Other Heinrich events are correlated with arid and coldPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 10 of 37

Table 3 Temperature and precipitation conditions during the LGM
Study site

Hemisphere

Zone

PeriodT

P

Citation(s)Global

NH/SH

0

30–15

Cold

0

Bush and Philander (1999)
Schmittner et al. (2011)

Global

NH/SH

0

30–15

0

Dry

Cragin et al. (1977)
Yung et al. (1996)
Steffensen (1997)
Li and Zhang (2020)
Bush and Philander (1998) Gasse (2000)

N. Hemisphere

NH

0

30–15

Cold

Dry

Bigelow et al. (2003)
Otto-Bliesner et al. (2006)
Yokoyama et al. (2000)

Tropics

NH/SH

L-L

30–15

Cold

0

Barker et al. (2005)
Annan and Hargreaves (2013)
Crowley (2000)
Ballantyne et al. (2005)

Tropics

NH/SH

L-L

30–15

0

Dry

Prentice et al. (2000)

Europe

NH

M-L

30–15

Cold

0

Strandberg et al. (2011)

Extra-tropics

NH

L-L

30–15

0

Wet

Clark and Mix (2002)

E. Africa

NH

L-L

30–15

0

Wet

Barker and Gasse (2003)

N. Africa

NH

M-L

30–15

0

Dry

Prentice et al. (2000)

S. Africa

SH

L-L

30–15

0

Dry

Patridge et al. (1997)
Simon et al. (2015)

N. Europe

NH

M-L

30–15

0

Dry

Strandberg et al. (2011)

S. Europe

NH

M-L

30–15

0

Wet

Strandberg et al. (2011); Kuhlemann et al. (2008)

Mediterranean

NH

M-L

30–15

0

Wet

Clark and Mix (2002)

SE. Siberia

NH

M-L

30–15

0

Dry

Osipov and Khlystov (2010)

E. & SW. Amazonia

SH

L-L

30–15

0

Dry

Absy et al. (1991) Sifeddine et al. (2001)

Greenland

NH

H-L

30–15

Cold

0

Johnsen et al. (1995)
Cuffey et al. (1995)
Miller et al. (2010)

Greenland

NH

H-L

30–15

0

Dry

Cuffey and Clow (1997)
Johnsen et al. (2001)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes. Period units are in ka BP and average age uncertainty
is ± 5 ka

climate at central China and even to Antarctica (Thompson 1991). Finally, they can also been detected over the
Indian Ocean (Bay of Bengal), linked to increased variability in the summer monsoon and drier conditions allover India (Colin et al. 1998; Wang et al. 2001). Similarly to the LGM or other glacial stages, there is strong
evidence that the decline in atmospheric/oceanic temperature results to the weakening or deceleration of thehydrological cycle and consequently to drier conditions(Mangerud et al. 2003; Grimm et al. 2006).
However, fluctuations to warm and wet conditions
have also been reported. Warmer SST has prevailed
over Southern California (Hendy and Kennett 2000),
while low isotopic values suggest an extremely wet
climate across the western USA between 40–30 ka
28.5–26.5 ka, and around 13 ka (Benson et al. 1996).
In addition, the δ18O records from the Owens Lake,
Great Basin (western United States) present overflow

conditions, which were caused by either high precipitation or enhanced ice melting (Gale 1914; Oster et al.
2014). The substantial growth in central Andean glaciers is an indication of increased precipitation acrosstropical South America during the Heinrich events 1
and 2 (Wang et al. 2004) and across northeast Brazil for
Heinrich events 1 to 5 (Smith and Rodbell 2010).
The studies on the Last Glacial abrupt climatic transitions are divided in warm (D–O) and cold (Heinrich)
events. All D–O events are associated with wet conditions, while the hydroclimatic shift for Heinrich events
is not so clear (Fig. 4 and Table 4). The cold transitions
appear to result to both dry and wet conditions, withthe dry conditions appearing more often (about a thirdof the studies). Thus, the hydrological cycle response
to Heinrich events appears more heterogeneous compared to the D–O events. Still the relatively low number of studies may affect these findings.

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 11 of 37

Fig. 4 Relationship of temperature and precipitation during the Last Glacial. A D–O events, B Heinrich events. The number of studies used for the
warm/cold or wet/dry conditions can be found in Table 4

3.2 Bølling–Allerød interstadial

During the final stages of the last glacial period, an abrupt
warm and moist period that occurred between 14.8 and12.85 ka BP (On et al. 2018). In some regions, the period
is divided into the¨ Bølling oscillation, with a peak closer

to 14.5 ka BP and duration around 1400 years, and theAllerød oscillation, with a peak around 13 ka BP and a
duration of 700 years (Seierstad et al. 2005). According tothe δ18O proxies of the GRIP ice core, the Bølling climate
was 1 °C colder than today, while Allerød was 5–12 °C

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 12 of 37

Table 4 Temperature and precipitation conditions during the Last Glacial
Study site

Hemisphere

Zone

Event

PeriodT

P

Citation(s)GlobalNH/SH

0

Heinrich

18–16.50

DryGates (1976)
Kwiecien et al. (2009)
Sun et al. (2019)

N. Hemisphere

NH

0

D–O

80–12

Warm

0

Martrat et al. (2004)

Tropics

NH/SH

0

Heinrich

70–14

0

Dry

Leuschner and Sirocko (2000)

EurasiaNH

M-L

Heinrich

70–14

Cold

Dry

Genty et al. (2003)
Benson et al. (1996)
Asmerom et al. (2010)

Asia

NH

M-L

D–O

80–12

0

Wet

Dung et al. (2020)
Cheng et al. (2016)
Kathayat et al. (2016)

Asia

NH

M-L

D–O

50–40

0

Wet

Wang et al. (2001)

N. America

NH

M-L

Heinrich

70–14

Cold

Dry

Genty et al. (2003)
Benson et al. (1996)
Asmerom et al. (2010)

S. America

SH

L-L

D–O

50–40

0

Wet

Peterson et al. (2000)

Antarctica

SH

H-L

Heinrich

70–14

Cold

Dry

Thompson (1991)

Tropical S. America

SH

M-L

Heinrich

15 and 220

Wet

Wang et al. (2004)

SW. Asia

NH

M-L

D–O

50–40

0

Wet

Anderson and Prell (1993)

W. Europe

NH

M-L

D–O

80–12

Warm

Wet

Genty et al. (2003)

N. EuropeNH

M-L

Heinrich

43–26

Cold

Dry

Mangerud et al. (2003)
Grimm et al. (2006)

Mediterranean

NH

M-L

Heinrich

70–14

Cold

0

Rohling et al. (1998)

India

NH

M-L

Heinrich

70–14

Cold

Dry

Colin et al. (1998) Wang et al. (2001)

Central China

NH

M-L

Heinrich

70–14

Cold

Dry

Thompson (1991)

Iberian Peninsula

NH

M-L

D–O

80–12

0

Wet

Nebout et al. (2002)
Budsky et al. (2019)

Greenland

NH

H-L

D–O

80–12

Warm

0

Johnsen et al. (1989) Lang et al. (1999)
Budsky et al. (2019)

Italy

NH

M-L

D–O

80–12

0

Wet

Alley and Clark (1999)

Israel

NH

M-L

D–O

80–12

0

Wet

Bar-Matthews et al. (2000)

NE. Brazil

SH

L-L

Heinrich

40–15

0

Wet

Smith and Rodbell (2010)

SW. USNH

M-L

Heinrich

70–14

0

Dry

Wagner et al. (2010)

W. US

NH

M-L

Heinrich

40–30

0

Wet

Benson et al. (1996)
Gale (1914)
Oster et al. (2014)

W. US

NH

M-L

Heinrich

28.5–26.5

0Wet

Benson et al. (1996)
Gale (1914)
Oster et al. (2014)

Florida, US

NH

M-L

Heinrich

70–14

Warm

Wet

Grimm et al. (2006)

Great Basin (Nevada, US)

NH

M-L

D–O

80–12

0

Wet

Benson et al. (1998)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes. Period units are in ka BP and average age uncertainty forthis period is ± 10 ka

colder (Johnsen et al. 1995). The lake sediments fromthe Lago di Origlio at Southern Swiss Alps suggest thatduring the Bølling–Allerød interstadial the temperature
increased about 2.5 to 3.2 °C (Samartin et al. 2012). Sediment analyses over the Aegean Sea and Lake Maliq show
an increased average annual temperature of about 10 °C
in the onset of the Bølling–Allerød interstadial, which
remained rather stable consequently (Bordon et al. 2009;
Kotthoff et al. 2011). Still, its onset is considered amongst

the most dramatic deglaciation events over the NH, possibly linked with the revival of the AMOC (Thiagarajan
et al. 2014).
The changes in Atlantic oceanic circulation intensified
the hydrological cycle over various regions across theglobe. One of the regions that were significantly affected
is the Mediterranean. Sediment analysis from Lake
Prespa (Greece) revealed enhanced humid conditions(Aufgebauer et al. 2012). Additionally, this increased

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

humid conditions were observed at Lake Maliq (Bordon
et al. 2009), Eastern Mediterranean (Bar-Matthews et al.
1999), and also Lago Grande di Monticchio (Italy) (Allenet al. 1999). At the same time, there was a widespreadincrease in both tropical and monsoon precipitation.
Significant increases are reported for equatorial Africa(Putnam and Broecker 2017; Tierney and deMenocal
2013), western Himalayas, Nepal and India (Sinha et al.
2005; Zech et al. 2014), and Northwest China (Zhou et al.
2001). Similar fluctuations in precipitation were observed
over Southern and Central America. Wet and warm conditions have been identified in lake sediments of Lagunade Los Anteojos (Venezuela) (Stansell et al. 2010), Pet´en
Itz´a (Guatemala) (Hodell et al. 2008), La Yeguada and El
Valle (Panama) (Bush et al. 1992) and Caribbean (Hughen
et al. 1996).
All the evidence suggest that the multi-centennial
increase in temperature was accompanied by an increasein precipitation too. In Fig. 5 and Table 5, we can see that
more than 90% of temperature records are confirming
warm conditions and about 85% of precipitation records
for wet conditions. Again, this abrupt transition suffersfrom a low number of studies, especially at larger spatial
scales (global, hemispheric, and continental).

Page 13 of 37

3.3 Younger Dryas

The Bølling–Allerød interstadial was followed by anothercool phase, the Younger Dryas event (from 13 to 11.7 ka
BP). An abrupt decline in temperature disrupted thegeneral warming trend that was driven by the increasing solar insolation (Dansgaard et al. 1989). Similarly to
LGM and Heinrich events, the drop in temperature wasaccompanied by generally dry conditions (Hodell et al.
2008; Mayewski et al. 1993; Mayewski and Bender 1995).
The Younger Dryas was mainly observed over the North
Atlantic region (Fairbanks 1990), but is also evident inpaleoclimatic records from all over the globe. However,the shift in the global climate was not homogeneous;contrary to the colder conditions of the high latitudes,
the tropics were characterized by comparatively warmer
conditions (Gagan et al. 2000). The temperature reconstructions of the Younger Dryas show a decline in temperature around 15 °C over central Greenland (Johnsen
et al. 1995), and a drop between 6 and 9 °C in the Norwegian Sea (Karpuz and Jansen 1992). There is no doubtthat Europe was substantially influenced by the Younger
Dryas event (Brauer et al. 2008; Rach et al. 2014). A 4 to
6 °C decrease over western Europe has been reported,reaching 6 to 7 °C over Poland (Go´slar et al. 1995). There

Fig. 5 Relationship of temperature and precipitation during the Bølling–Allerød interstadial. The number of studies used for the warm/cold or wet/dry conditions can be found in Table 5

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 14 of 37

Table 5 Temperature and precipitation conditions during the Bølling–Allerød interstadial
Study site

Hemisphere

Zone

PeriodT

P

Citation(s)N. Hemisphere

NH

L-L

13–12

0

Wet

Knox and Wright (1983)
Maher (2008)
Pausata et al. (2020)

S. & C. America

NH

M-L

14.8–12.85

Warm

Wet

Stansell et al. (2010)
Hodell et al. (2008)
Bush et al. (1992)
Hughen et al. (1996)

Equatorial Africa

NH

M-L

14.8–12.85

0Wet

Putnam and Broecker (2017) Tierneyand deMenocal (2013)

Arctic

NH

H-L

14.5

Cold

0

Johnsen et al. (1995)

E. Mediterranean

NH

M-L

14.8–12.85

0Dry

Bar-Matthews et al. (1999)

GreeceNH

M-L

14.8–12.85

0DryAufgebauer et al. (2012)

Italy

NH

M-L

14.8–12.85

0Dry

Allen et al. (1999)

W. USA

NH

M-L

13

0

Wet

Benson et al. (1996)
Gale (1914)
Oster et al. (2014)

Aegean Sea

NH

M-L

14.8–12.85

Warm

0

Bordon et al. (2009) Kotthoff et al. (2011)

S. Swiss Alps

NH

M-L

14.8–12.85

Warm

0

Samartin et al. (2012)

W. Himalayas

NH

M-L

14.8–12.85

0WetSinha et al. (2005)

NH

M-L

14.8–12.85

Warm

0

Bordon et al. (2009) Kotthoff et al. (2011)(Nepal & India)
Lake Maliq, Albania

Zech et al. (2014) Zhou et al. (2001)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes. Period units are in ka BP and average age uncertainty
is ± 1 ka

is also evidence of re-extension of the North European
ice sheet (Aufgebauer et al. 2012). Consequently, thisprocess led to the southerly flow of dry and cold northern
air masses towards the Mediterranean area (Bordon et al.
2009), which led to colder temperatures in the Aegean
Sea (Kotthoff et al. 2011). On the other hand, a pollen
record from east Beringia revealed a more constrained
drop in temperature, estimated at 1.5 °C (Fritz et al.
2012), which is in agreement with evidence that several
coastal areas near western Novaya Zemlya (Russia) were
ice-free (Serebryanny et al. 1998).
In terms of hydroclimate, various regions of the NH
have experienced drier conditions during the Younger
Dryas (Dahl and Nesje 1992; Fawcett et al. 1997; Hughen
et al. 2000; Starkel 1991; Velichko et al. 2002). However,there are many regions that did not maintain a stable
cold and dry regime, but instead the cold climate wascoupled by centennial oscillations between dry and wetphases (Wang et al. 2018). For instance, the increasein the hydrogen isotope values at about 12 ka BP and
12.2 ka BP suggesting wetter and warmer phases overwestern Europe (Rach et al. 2014). This is also evident
in Central Europe, where paleoclimatic records (Magny
2001) and certain periglacial characteristics (Kaiser andClausen 2005) suggest wet conditions during the Younger
Dryas (Weber et al. 2011), especially during winter (Isarin and Bohncke 1999). Other evidence of precipitationcomparative to the present has been recorded in Poland(Prosna River, about 30% higher) (Rotnicki 1991), Netherlands (Bos et al. 2006), and Scotland highlands (Lukas
and Bradwell 2010). A multi-proxy reconstruction from
central Poland for the Younger Dryas reports two phases(Pawl owski et al. 2015). The first (12.5–12 ka BP) was
marked by a decrease in precipitation and temperaturesduring winter, but a rise in summer precipitation. The
second (12–11.5 ka BP) shows increased winter and summer temperature with increased annual precipitation. In
the southern Europe, pollen records from Mediterranean
show increased precipitation during the whole deglaciation phase (18–10 ka), without any influence by Younger
Dryas (Paterne et al. 1999).
A zonal gradient in precipitation response appears inNorth America. Drier conditions dominate the northern parts of the continent (Carlson 2013; Dorale et al.
2010), transitioning to considerably wetter conditions as
we move southwards (Grimm et al. 2006; Voelker et al.
2015). There, the precipitation levels have been estimatedat about 15% higher values than today, probably due toincreased southern atmospheric moisture flow (Renssen et al. 2018). Wetter phases over central and southern
North America are further supported by various proxyevidence in plant-macrofossil and palynology studies
over Florida, and speleothems from New Mexico (Polyak
et al. 2004) and Arizona (Wagner et al. 2010). Climate

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

model simulations also present a warmer climate withincreased precipitation in the central regions of North
America during the Younger Dryas when compared to
the Bølling–Allerød period (Renssen 2020).
In the SH, there are conflicting results. Some studies
provide evidence for enhanced precipitation and conditions similar to the Heinrich Events (Arz et al. 1998). For
instance, the analysis lake sediments from Lake Titicaca(Bolivia, Peru) demonstrates the overflowing of the lake
and thus higher precipitation between 13 and 11.5 ka BP(Baker et al. 2001). On the other hand, the lake sedimentrecords at the lake Laguna de Los Anteojos (Venezuela)
present a transition to an intense cold and dry regime
during the Younger Dryas (Stansell et al. 2010). Thisis further supported by a significant drop in the Amazon River discharge that is probably a result of reduced
monsoon precipitation over the lowland tropical SouthAmerica (Maslin and Burns 2000). Moreover, arid conditions are reported across the northern tropical Andes
and wetter conditions over the southern tropical Andes(Stansell et al. 2010).
With 80% of the studies revealing a transition to cold
conditions, there is little doubt about the temperature conditions of the Younger Dryas (Fig. 6). The same

Page 15 of 37

cannot be said about the hydroclimatic regime, where
studies remain split almost in the half. About 60% of therecords suggest wet conditions, while the remaining 40%
present a drier climate. The majority of the studies atlarger spatial scales (global and N. Hemisphere) point todry conditions, while wet conditions are more frequentin the finer scales.

4 Climatic fluctuations in the Holocene
4.1 The 8.2 ka event

The’8.2 ka cold event’ is another abrupt climatic eventthat was experienced across the entire globe originatingfrom the North Atlantic region (Alley and Ágústsdóttir
2005). As the name implies, it occurred around 8.2 ka BP
and lasted for 160.5 ± 5.5 years, with the coldest period
spanning 69 ± 2 years (Thomas et al. 2007). Other estimates suggest a duration between 150 and 200 years (Von
Grafenstein et al. 1998; Snowball et al. 2002). The available proxy records show an abrupt cooling up to 6 ± 2 °C
(Allen et al. 2007; Alley et al. 1997; Dansgaard et al. 1993),
resulting to a global decrease by 0.9–1.8 °C (Heikkilä and
Seppä 2010). Greenland is one of the regions with themost intense drops, about 3 to 8 °C (Alley et al. 1997; Von
Grafenstein et al. 1998), as well as, enhanced windy andFig. 6 Relationship of temperature and precipitation during the Younger Dryas. The number of studies used for the warm/cold or wet/dry
conditions can be found in Table 6

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 16 of 37

Table 6 Temperature and precipitation conditions during the Younger Dryas
Study site

Hemisphere

Zone

PeriodT

P

Citation(s)GlobalNH/SH0

13–11.7

Cold

Dry

Hodell et al. (2008)
Mayewski et al. (1993)
Mayewski and Bender (1995) Fairbanks (1990)

N. Hemisphere

NH

H-L

13–11.7

Cold

0

Gagan et al. (2000)

N. HemisphereNH0

13–11.7

0DryDahl and Nesje (1992)
Fawcett et al. (1997)
Hughen et al. (2000), Starkel (1991)
Velichko et al. (2002)

Tropics

NH/SH

L-L

13–11.7Warm

0Gagan et al. (2000)

Norwegian Sea

NH

H-L

13–11.7

Cold0

Karpuz and Jansen (1992)

N. America

NH

M-L

13–11.70

DryCarlson (2013)

NH

M-L

13–11.7

0

Wet

Grimm et al. (2006)

NH

M-L

13–11.7

0

Wet

Polyak et al. (2004)

NH

M-L

13–11.7

Warm

Wet

Renssen (2020)(Northern part)N. America

Dorale et al. (2010)(Southern part)
Central & Southern

Voelker et al. (2015) Renssen et al. (2001)

N. America
N. America (Central)

Wagner et al. (2010)

W. EuropeNH

M-L

13–11.7

Cold0

Go´slar et al. (1995)

W. EuropeNH

M-L

12–12.2

Warm

Wet

Rach et al. (2014)

Central Europe

NH

M-L

13–11.7

0

Wet

Magny (2001)
Kaiser and Clausen (2005) Weber et al. (2011)
Isarin and Bohncke (1999)

S. Europe

NH

M-L

18–10

0

Wet

Paterne et al. (1999)

N. Tropical Andes

SH

M-L

13–11.70

DryStansell et al. (2010)

S. Tropical Andes

SH

M-L

13–11.7

0

Wet

Stansell et al. (2010)

Tropical S. America

SH

L-L

13–11.70Dry

Maslin and Burns (2000)

Bolivia & Peru

SH

L-L

13–11.7

0

Wet

Baker et al. (2001)

PolandNH

M-L

13–11.7

Cold0

Go´slar et al. (1995)

PolandNH

M-L

13–11.7

0Wet

Rotnicki (1991)

Netherlands

NH

M-L

13–11.7

0Wet

Bos et al. (2006)

Venezuela

NH

L-L

13–11.7

Cold

Dry

Stansell et al. (2010)

E. Beringia

NH

M-L

13–11.7

Cold

0Fritz et al. (2012)

N. Scotland

NH

M-L

13–11.7

0

Wet

Lukas and Bradwell (2010)

Aegean Sea

NH

M-L

13–11.7

Cold

0

Kotthoff et al. (2011)

Central Greenland

NH

H-L

13–11.7

Cold

0

Johnsen et al. (1995)

NH: Northern Hemisphere, SH: Southern Hemisphere, H-L: High-latitudes, M-L: Mid-latitudes, L-L: Low-latitudes. Period units are in ka BP and average age uncertainty
is ± 0.5 ka

dry conditions over most parts of the NH (T¨ornqvist
et al. 2004) at a time when the climatic conditions weresimilar as of today (Alley et al. 1997).
The areas with the most rapid transition were widespread across the entire Baltic Sea basin (Borzenkova
et al. 2015), the western Europe (Davis et al. 2003), andthe regions affected by the NAO, in particular (Sepp¨aet al. 2008). The latter experienced a decline between
1.5 and 3 °C, as both land and marine records suggest(Klitgaard-Kristensen et al. 1998; Bond et al. 1997; Von
Grafenstein et al. 1998). The above results are in good
agreement with model simulations. The models present
cooling around 2 to 5 °C over Greenland (Gasse 2000),

2.5 °C at the lake Annecy (France) (Magny et al. 2003), 1
to 2 °C over northwestern Europe (Renssen et al. 2001),
as well as approximately 2 °C over Germany and theNorth Sea (Klitgaard-Kristensen et al. 1998).
Some model simulations also suggest a 30% drop in
precipitation (Gasse 2000). This is in good terms with the
dry conditions which have been generally observed over
the NH (Clarke et al. 2004; Alley et al. 1997), particularly in the wintertime (Alley and Ágústsdóttir 2005). In
Europe, the transition to dryer conditions was observed
to latitudes over 50°N, as well as a significant part of theMediterranean, including Spain, Northern Africa, andItaly (Magny and B´egeot 2004). On the other hand,Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

during the 8.2 event, a worldwide snowfall increase wasobserved (Borzenkova et al. 2015). This could explain thelake-level rises in many European palaeoclimate records,
related to higher runoff (Magny 1992). The lake-level
rise becomes more evident over the central Alps (Switzerland, France, and northern Italy) (Magny and B´egeot2004).
All the evidence suggest that the 8.2 event was characterized by colder and drier climate conditions (Fig. 7),with only two studies presenting wet conditions over
the Alps (Table 7). Even though there is good agreement
between the records at both coarse and fine spatial scales,
we cannot rule out though a small-sample bias in thisconclusion due to the limited number of studies describing the precipitation of this cold period.
4.2 Medieval climate anomaly

The Medieval Climate Anomaly (MCA), also known
as Medieval Warm Period, is the most recent period ofabrupt warming, with onset around 800–1000 CE andtermination at 1300–1400 CE (Hughes and Diaz 1994).
It affected mostly Europe and parts of North America,
which mainly experienced warmer than average conditions (Lamb 1965). The centennial-scale patterns ofPage 17 of 37

spatiotemporal temperature reconstructions suggest
widespread warm and arid conditions over the NH with asimilar geographic extent and magnitude as in the twentieth century mean (Ljungqvist et al. 2016). Between
1200 and 1300 CE, the temperature was similar to thepresent over northwestern Europe (Guiot 1992). In addition, a temperature reconstruction across the Alps suggests that in the twelfth century the temperature was
0.3 °C higher than today (Trachsel et al. 2012). In North
America, there is conflicting evidence about the increasemagnitude. Viau et al. (2012) demonstrated that there
was a 0.5 degree increase, which resulted to cooler thanthe present conditions, whereas Woodhouse et al. (2010)
report temperatures of about 1 °C higher than today.
There is also evidence of high temperatures over China(Yang et al. 2002), South Atlantic (Jones and Mann 2004),
and Northern Pacific (Mann et al. 2009). Even though theextent of temperature increases during MCA remains
under investigation, there is general agreement that there
has been a clear signal of the increase at least in the NH.
The hydroclimatic response, though, was not so homogeneous. Substantial precipitation deficiencies wereobserved in northern Europe (Cook et al. 2015) and East
Africa (Verschuren et al. 2000). In addition, model andFig. 7 Relationship of temperature and precipitation during the 8.2 ka cold event. The number of studies used for the warm/cold or wet/dry
conditions can be found in Table 7

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 18 of 37

Table 7 Temperature and precipitation conditions during the 8.2 event
Study site

Hemisphere

Zone

Time

T

P

Citation(s)Global

NH/SH

0

8.2

Cold

0

Alley and Ágústsdóttir (2005)
Thomas et al. (2007)
Heikkilä and Seppä (2010)

N. Hemisphere

NH

0

8.2

Cold

Dry

T¨ornqvist et al. (2004)
Clarke et al. (2004) Alley et al. (1997)

North Sea

NH

M-L

8.2

Cold

0

Klitgaard-Kristensen et al. (1998)

N. Africa

NH

M-L

8.2

Cold

Dry

Magny and B´egeot (2004)

NW. Europe

NH

M-L

8.2

Cold

0

Renssen et al. (2001)

Greenland

NH

H-L

8.2

Cold

0

Alley et al. (1997)
Von Grafenstein et al. (1998) Gasse (2000)

Germany

NH

M-L

8.2

Cold

0

Klitgaard-Kristensen et al. (1998)

France

NH

M-L

8.2

Cold

0

Magny et al. (2003)

Spain

NH

M-L

8.2

Cold

Dry

Magny and B´egeot (2004)

Italy

NH

M-L

8.2

Cold

Dry

Magny and B´egeot (2004)

Central Alps

NH

M-L

8.2

0

Wet

Magny (1992); Magny and B´egeot (2004)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes. Period units are in ka BP

paleoclimatic records show that western North America
experienced persistent and extensive aridity from 900 to
1300 CE (Woodhouse et al. 2010). On the other hand,anomalously lower δ18C values in bristlecone pine fromthe White Mountains, California, highlight a wet period
from 1080 to 1129 CE (Hughes and Diaz 1994). Similarly,
Mauquoy et al. (2004) also suggested the times from 1030
to 1100 CE was a wet period for western North America,
while there is also evidence of higher lake levels, or freshwater availability, over the Arizona monsoon-influenced
area from 700 to 1350 CE (Hughes and Diaz 1994). However, the occurrence of increased aridity over most of
the areas of western North America was also evident intree-ring records between 650 and 1050 CE (Parish et al.
2020) and from 900 to 1300 CE (Cook et al. 2007).
In Asia, dry climatic conditions prevailed, mainly
linked with atmospheric circulation (Chen et al. 2015).
This in good terms with the proxy analysis over southern
China indicates comparatively weak monsoonal precipitation over most of the regions (Chen et al. 2015) and theperiods of extensive aridity from 1140 to 1220 and 1420
to 1490 (Li et al. 1987). On the contrary, pollen estimatesfrom Maili pond at northeast China reveal wet conditions (from 950 to 1290 CE), suggesting an increase in
the East Asian summer monsoon (EASM) during thisperiod (Ren 1998). Additionally, the decades between
1230–1250 CE, and 1380–1410 CE show intensificationof the South-Asian monsoon resulting to wet conditions(Li et al. 1987). Most of the proxy records suggest precipitation decrease over the EASM region after the termination of MCA around 1300 CE (Lan et al. 2020).
South America also experienced a highly variable climate during MCA. Perhaps this is due to a humidity

dipole between the southern and northern Amazon Basin(Marengo 2004). This humidity dipole could suggest an
enhanced land–ocean temperature gradient or north–south migration of the ITCZ, driven by seasonal variationin the distribution of insolation (Wright et al. 2017). Consequently, the wetter phase over the northeast area wassynchronous with the drier phase over Southern Amazonia (Thompson et al. 2013). For instance, a marine sediment core at Peru ­(12◦S) shows intense aridity between
800 to 1250 CE (Rein et al. 2004), while a titanium (Ti)record from the Cariaco Basin (Venezuela) indicates wetter conditions between 950 to 1450 CE (Haug et al. 2001).
Additionally, the assessment of lake sediment oxygenisotopes (δ18O) at the Central Peruvian Andes presents
higher values from 900 to 1100 CE, implying a weakened
South American Summer Monsoon and a prolonged
period of aridity (Bird et al. 2011). On the other hand,
wet conditions prevailed in Central America, as indicate
by the lower values of oxygen isotope in sediments from
Nicaragua from 950 to 1250 CE (Stansell et al. 2013).
Another feature of MCA is the emergence of simultaneous mega-droughts in various regions of the globe(Stager et al. 2005). The main region affected of these
multi-decadal droughts can be found at North America(Cook et al. 2014). There, two prolonged drought eventswith an approximately 90 years time span have beenrecorded over North America. The first event occurred
between 1197 and 1289 CE, while the second event
occurred between 1486 and 1581 CE (Parish et al. 2020).
Other shorter events have been also detected, presenting higher severity, though, such as the mega-droughtfrom 1140 to 1162 CE or the one between 1150 to 1159
CE (Cook et al. 2007). In Europe, the multi-decadalPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

reconstruction over the Sierra Nevada (Spain) highlights
four multi-decadal droughts that prevailed during the
MCA (800–859 CE, 1020–1070 CE, 1197–1217 CE, and
1249–1365 CE) (Graumlich 1993).
The main hypothesis about the driver of the enhancedhydroclimatic variability of MCA is the positive state of
NAO, which persisted at centennial time scale (Trouet
et al. 2009). The result was a northeastward shift of thecyclonic storm tracks, and consequently the transport ofatmospheric moisture to higher latitudes (Solomon et al.
2007). The spatial hydroclimatic variability is also evident
in finer scales. A typical case is highlighted over the Iberian Peninsula. There, a climate reconstruction shows
warmer and humid conditions across the northwest
regions, while the rest of the peninsula shows warm andarid conditions (Moreno et al. 2012). Similar patterns canbe seen in tree-ring records over Morocco, where some
unusually frequent wet years occurred from 1250 to 1300
CE (Till et al. 1990).
During the MCA, all the studies analyzed clearly suggest a warmer climate (Fig. 8 and Table 8). However, contrary to the other warm periods presented in this study,
the multi-centennial warming was coupled with dry
conditions. About two thirds of the records indicate aPage 19 of 37

transition to a dry climate, which might be seen as contradiction to the prevailing theory of water cycle intensification and will be discussed in detail below.
4.3 Little Ice Age

The Little Ice Age (LIA) is the most recent shift to colder
conditions. It lasted from 1350–1450 CE to 1900 CE(Mayewski and Bender 1995), and the global temperature
was 0.5 to 1.5 °C lower than the twentieth century average(Crowley and North 1991; Graumlich 1993; Mann et al.
1998; Christiansen and Ljungqvist 2012; Schneider et al.
2015). Trachsel et al. (2012) have reported that during
fourteenth, late sixteenth, and seventeenth century, theglobal temperature was a 1 °C lower than the twentiethcentury average. The NH experienced the most substantial decrease (about 0.9 °C lower) from 1570 to 1730 CE(Bradley and Jonest 1993), whereas in Europe LIA peakedin 1650–1750 CE (Bond et al. 2001). In SH, paleoclimaticoceanic records show an average cooling of 1.6 °C (± 1.4)
compared to the last 150 years (Rhodes et al. 2012). Ice
core analysis near the Ross Sea (Antarctica) shows colder
conditions of 2 °C in surface temperature, as well as lower
SST over the Southern Ocean coupled by enhanced sea
ice extent during the LIA (Bertler et al. 2011). In general,

Fig. 8 Relationship of temperature and precipitation during the MCA. The number of studies used for the warm/cold or wet/dry conditions can befound in Table 8

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 20 of 37

Table 8 Temperature and precipitation conditions during the MCA
Study site

Hemisphere

Zone

Time

T

P

Citation(s)N. HemisphereNH0800–1400

Warm

Dry

Ljungqvist et al. (2016)

Asia

NH

M-L

900–13000

DryChen et al. (2015)

N. America

NH

M-L

1197–1289

0

Dry

Cook et al. (2014)

0

Dry

Cook et al. (2007)

And 1486–1581
N. America

NH

M-L

1140–1162

Parish et al. (2020)

N. America

NH

M-L

800–1400Warm

0Woodhouse et al. (2010)

N. America

NH

M-L

800–1400Warm

0Lamb (1965)

EuropeNH

M-L

800–1400Warm

0Lamb (1965)

N. EuropeNH

M-L

800–1400

0

Dry

Cook et al. (2015)

S. AsiaNH

M-L

1230–1250
and 1380–1410

0Wet

Li et al. (1987)

E. Asia

NH

M-L

1300

0

Dry

Lan et al. (2020)

NW. Europe

NH

M-L

1200–1300

Warm

0

Guiot (1992)

S. America (NE area)

SH

L-L

800–1400

0

Wet

Thompson et al. (2013)

Western N. America

NH

M-L

900–1300

Warm

Dry

Woodhouse et al. (2010)

Western N. America

NH

M-L

1030–1100

0

Wet

Mauquoy et al. (2004)

Western N. America

NH

M-L

650–1050

0

Dry

Parish et al. (2020)

Western N. America

NH

M-L

900–1300

0

Dry

Cook et al. (2007)

E. AfricaNH

M-L

800–1400

0

Dry

Verschuren et al. (2000)

S. Atlantic

SH

M-L

800–1400Warm

0Jones and Mann (2004)

N. Pacific

NH

M-L

800–1400

Warm

0Mann et al. (2009)

China

NH

M-L

800–1400Warm

0Yang et al. (2002)

S. ChinaNH

M-L

800–14000

Dry

Chen et al. (2015)

S. China

NH

M-L

1140–1220
and 1420–1490

0Dry

Li et al. (1987)

NE. ChinaNH

M-L

950–1290

0Wet

Ren (1998)

Southern Amazonia

SH

L-L

800–14000Dry

Thompson et al. (2013)

Peru

SH

L-L

800–1250

0Dry

Rein et al. (2004)

MoroccoNH

M-L

1250–1300

0Wet

Till et al. (1990)

Central America

NH

M-L

950–1250

0

Wet

Stansell et al. (2013)

Spain (Sierra Nevada)

NH

M-L

800–859,
1020–1070,
1197–1217,
And 1249–1365

0

Dry

Graumlich (1993)

Venezuela

NH

L-L

950–1450

0

Wet

Haug et al. (2001)

Arizona, USA

NH

M-L

700–1350

0Wet

Hughes and Diaz (1994)

California, USA

NH

M-L

1080–1129

0

Wet

Hughes and Diaz (1994)

Iberian Peninsula

NH

M-L

800–1400

Warm

Dry

Moreno et al. (2012)

Central Peruvian Andes

SH

L-L

900–11000Dry

Bird et al. (2011)

Alps

NH

M-L

twelfth century

Warm

0

Trachsel et al. (2012)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes. Period units are in AD and average age uncertainty for
the MCA is ± 100 years

the lowest temperatures were observed in the period of
1680 to 1730 CE, for both Hemispheres (Stuiver et al.
1995).
The LIA has been compared to the abrupt changes
that occurred in the last glacial stage (Bond et al. 1999),
such as the D–O events (Broecker 2000). However, eventhough LIA affected the whole globe, this did not happen

simultaneously. Most local or regional paleoclimaticreconstructions show unusually cold phases from 1580
to 1880 CE, interrupted by decades of warmer conditions(Ahmed et al. 2013). Similarly to MCA, the main hypothesis for the spatio-temporal variability lies in the changesof atmospheric circulation (Zhang et al. 2021a). Compared to the current patterns of atmospheric circulation,

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

LIA experienced stronger meridional transport (Lamb2013). This was observed over the North Atlantic andpolar South Pacific at the beginning of the LIA, evident
in ice cores from central Greenland, Siple Dome, andWest Antarctica (Kreutz et al. 1997). The colder and drier
conditions that prevailed were a result of the enhanced
atmospheric circulation, as reported in numerous paleoclimatic records in the NH and the Equator (Thompson
et al. 1995; O’Brien et al. 1995). Additionally, numericalmodel experiments have identified sea ice-ocean–atmosphere (Zhong et al. 2011) and volcanic feedbacks (Miller
et al. 2012) as a factor that triggered the LIA cooling over
the North Atlantic and Europe.
The drop of temperature was coupled to wet conditions over the most of the European territory (Luoto andNevalainen 2018; Brönnimann et al. 2019). Both the speleothem record from Scotland (Proctor et al. 2000) anda reconstruction from England-Wales (Lamb 1965) arenotably similar, showing a 10% decrease in the precipitation (for September to June) from the late thirteenth tothe mid fourteenth century, and a constant drop fromthe mid of sixteenth to late eighteenth century. Precipitation reconstructions from southern Moravia (Czech
Republic) show that the highest precipitation occurred
between 1670 and 1710 CE, succeeding a period withlow precipitation (Br´azdil et al. 2002). Proxy estimates ofseasonal precipitation over Europe exhibit increased winter (DJF) precipitation during the beginning of the eighteenth century (Pauling et al. 2006), which is attributed toa significant increase in winter temperatures (Nesje et al.
2008). This is in good agreement with the abrupt increasein floods reported from 1760 to 1800 CE over various
locations (Blöschl et al. 2020). Other similar periods are
1560–1580 and 1840–1870, when the climate conditions
were abruptly shifted to a warmer phase (Glaser et al.
2010) and consequently increasing precipitation and/or
snowmelt (Br´azdil et al. 1999). This is particularly true
for the end of the LIA, when there has been a monotonic
increase towards more humid conditions (Cook et al.
2015; Markonis et al. 2018).
Over the North American continent, there is evidenceof strong spatiotemporal heterogeneity in the observedchanges. In general, wetter conditions were observed in
the central regions compared to the present, while drier
conditions prevailed over both the West and East Coast(Ladd et al. 2018). In most of the wet periods, precipitation increased during the winter (Parish et al. 2020),
lasted for a couple of decades and were succeeded bylong dry intervals (Meko 1992). For instance, a precipitation reconstruction at Banff, Alberta (Canada) showshigher precipitation from 1515 to 1550 CE, 1585 to 1610
CE, 1660 to 1680 CE, and during the 1880s, while 1950 to1970s exhibit both enhanced precipitation and decreased

Page 21 of 37

summer temperatures (Luckman 2000). On the other
hand, the spatio-temporal drought and precipitation
records over North America suggest a widespread limitation in moisture availability during the late sixteenth
century while relative abundance during the early seventeenth century (Cook et al. 1997; Bradley et al. 2003; Matthews and Briffa 2005).
Various changes are reported in the rest of the world,
related to the fluctuations of atmospheric circulation.
Sediment records from the northeastern Arabian Sea
show a weakening of Indian summer monsoon from 1450
to 1750 CE and consequently a shift to drier conditions(Agnihotri et al. 2002). Northern China also faced a moderately weak monsoon (Chen et al. 2015). The lakelevels
and diatom estimates over Africa (Street-Perrott andPerrott 1990), and dust records in an equatorial ice core
(Thompson et al. 1995) also display increased aridity. The
paleoclimate records from the Argentina show during
about 1800 and 1930 as the wet period (Mauquoy et al.
2004). However, the isotope (increased values) evidence
from Central America suggests the persistence of drier
conditions during most of the LIA (Stansell et al. 2013).
Additionally, the tree-ring analysis from southern SouthAmerica indicates cold-dry/drought phase between 1280
and 1450, 1550 and 1670, and 1780 to 1830 CE; while thewarm-wet/high-rainfall phases from 1220 to 1280, 1450
to 1550, 1720 to 1780, and 1830 to 1905 CE (Villalba1994).
Increased precipitation was also observed in various
regions across the world. Low concentration of microparticles in ice core records from the Antarctic Peninsula indicates likely higher precipitation and intense
cyclonic activity (Rogers 1983). The enhanced meridional circulation has been expected to influence the mid
and low latitude circulation, resulting to a shift of thewesterlies belt and increased precipitation over the
Patagonia and California around 1400 CE (Stine 1994).
Additionally, the arid central Asia region is showing relatively wet conditions, and pluvial conditions prevailed
over southern China (Chen et al. 2015). The wet conditions were often succeeded by arid conditions, resulting to 18 extreme flood and 16 drought events during
the LIA in China (Zheng et al. 2006). Similarly, sediment geochemistry from a subalpine lake at northern
Taiwan indicated four pluvials (1660 CE, 1730, 1820,
and about 1920) (Wang et al. 2013; Zhao et al. 2018).
In South America, the oxygen isotope (δ18O) estimatesof a speleothem record at northeastern Peru reportenhanced variability in precipitation, with annual precipitation being 10% higher than today from fifteenth toeighteenth century (Reuter et al. 2009). This is in good
agreement with the results of an oxygen isotopes (δ18O)
analysis at the Central Peruvian Andes lake, showing aPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

prolonged regionally synchronous intensification in theSouth American Summer Monsoon (Bird et al. 2011).
Similar conclusions were drawn in the study of Polissar et al. (2006) about the growth of glaciers at the highelevations over the Venezuelan Andes, which can beinterpreted as evidence of higher precipitation.
The analysis of the corresponding literature advocates that LIA is not homogeneous event in space
and time. There are approximately 25% of studies thatreveal some region and/or period of warm conditions(Fig. 9). This is due to the availability of higher resolution reconstructions, which can detect shorter warmer
periods within the prevailing cold conditions, suchas for example the 1560–1580 and 1840–1870 warm
intervals over Europe (Br´azdil et al. 1999; Glaser et al.
2010). In addition, there are numerous locations withcold and wet conditions, resulting to a sum of 60% of
studies presenting a wet LIA, and 40% of records suggesting otherwise. Similarly to MCA, this is a reversed
relationship between temperature and precipitationcompared to the other periods studied. A possible
explanation for this outcome could lie to the fact thatthe majority of the studies come from Europe, amplifying the wet signal (Table 9).

Page 22 of 37

5 Insights from the past
Although our literature review study focuses in providing the empirical evidence of past hydroclimatic changes,
in this last Section we will briefly discuss some plausible
explanations for our findings. Perhaps the most striking
result is that even during the highest temperature deviations amongst the ones we examined, the hydrologicalcycle fluctuated within a reasonable range. No extreme
cases of global long-term aridity or humidity have been
imprinted in the paleoclimatic records. On the contrary, most climatic shifts present substantial spatial
heterogeneity regardless of their time scale. Of course,
different physical mechanisms will drive hydroclimatic
variability in different spatio-temporal scales. Due to thelarge uncertainties involved and the scarcity of the data
records, it is rather questionable if the exact processes
could be described, though. What could be more pragmatic is to distinguish the impact of the thermodynamic
and dynamic component.
Higher temperatures appear more strongly related towet conditions than lower temperatures to dry (Table 10).
Out of the five warm periods studied, four present a distinct warm-and-wet signal and only during the MCA thedry conditions prevailed. On the other hand, only two

Fig. 9 Relationship of temperature and precipitation during the LIA. The number of studies used for the warm/cold or wet/dry conditions can befound in Table 9

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 23 of 37

Table 9 Temperature and precipitation conditions during the LIA
Study site

Hemisphere Zone TimeT

P

Citation(s)GlobalNH/SH

01350–1900 CE

Cold

0

Crowley and North (1991)
Graumlich (1993)
Mann et al. (1998)
Christiansen and Ljungqvist (2012) Schneider
et al. (2015)

GlobalNH/SH

0

14th, Late sixteenth and seventeenth
centuryCold0

Trachsel et al. (2012)

N. HemisphereNH0

1570–1730 CE

Cold

0

Bradley and Jonest (1993)

S. HemisphereSH01350–1900 CE

Cold

0

Rhodes et al. (2012)

AfricaNH

L-L

1350–1900 CE

0

Dry

Street-Perrott and Perrott (1990)

Antarctica

SH

H-L

1350–1900 CE

Cold

0

Bertler et al. (2011)

EuropeNH

M-L

1650–1750 CE

Warm 0Bond et al. (2001)

EuropeNH

M-L

Early eighteenth century

0EuropeNH

M-L

Early eighteenth century

Warm 0EuropeNH

M-L

1760–1800 CE

0EuropeNH

M-L

1560–1580 CE

Warm Wet Glaser et al. (2010) Br´azdil et al. (1999)

EuropeNH

M-L

1840–1870 CE

Warm Wet Glaser et al. (2010) Br´azdil et al. (1999)

N. America

NH

M-L

Late ­16th

0

Wet Pauling et al. (2006)
Nesje et al. (2008)

Wet Blöschl et al. (2020)

Dry

and Early seventeenth century

Cook et al. (1997)
Bradley et al. (2003) Matthews and Briffa(2005)

Central Asia

NH

M-L

1350–1900 CE

0

Wet Chen et al. (2015)

Central America

NH

M-L

1350–1900 CE

0Dry

N. America(Central regions)

NH

M-L

1350–1900 CE

0

Wet Ladd et al. (2018)

N. America(W & E Coast)

NH

M-L

1350–1900 CE

0

Dry

Ladd et al. (2018)

Southern S. America SH

M-L

1280–1450 CE,
1550–1670 CE,
and 1780–1830 CE

Cold

Dry

Villalba (1994)

Southern S. America SH

M-L

1220–1280 CE
1450–1550 CE,
1720–1780 CE,
and 1830–1905 CE

Warm Wet Villalba (1994)

Antarctic Pen

SH

H-L

1350–1900 CE

0

Wet Rogers (1983)

Arabian Sea

NH

L-L

1450–1750 CE

0

Dry

Argentina

SH

M-L

1800 and 1930

0

Wet Mauquoy et al. (2004)

CanadaNH

M-L

1515–1550 CE,
1585–1610 CE,
1660–1680 CE,
and 1880s

0

Wet Luckman (2000)

Czech Republic

NH

M-L

1670–1710 CE

0

Wet Br´azdil et al. (2002)

Venezuelan Andes

NH

L-L

1400–1700 CE

0

Wet Polissar et al. (2006)th

Stansell et al. (2013)

Agnihotri et al. (2002)

ScotlandNH

M-L

Late ­13 to mid fourteenth century

0

Wet Proctor et al. (2000)

Peru

SH

L-L

1400–1700 CE

0

Wet Reuter et al. (2009)
Bird et al. (2011)

California, USA

NH

M-L

1400 CE0

Wet Stine (1994)

Wales, UKNH

M-L

Late 13th mid fourteenth century

0

Wet Lamb (1965)

EnglandNH

M-L

Late 13th mid fourteenth century

0

Wet Lamb (1965)

Patagonia

NH

M-L

1400 CE0

Wet Stine (1994)

N. China

NH

M-L

1450–1750 CE

0Dry

N. TaiwanNH

M-L

1660, 1730, and 1820 CE

0

Wet Wang et al. (2013)

Chen et al. (2015)
Zhao et al. (2018)

NH Northern Hemisphere, SH Southern Hemisphere, H-L High-latitudes, M-L Mid-latitudes, L-L Low-latitudes and average age uncertainty for the LIA is ± 50 years

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 24 of 37

Table 10 Number of studies per period and conditions. The Hydroclimate column describes the most common condition. Average
duration is estimated in thousand years
PeriodDuration (ky)Studies

Warm

Cold

Dry

Wet

Hydroclimate

MMCO

2500

40

49

0

22

39

Warm and wet

MIS-5e

14

32

12

2

7

21

Warm and wet

LGM

15

30

0

13

20

5

Cold and dry

D–O events

0.8

15

5

0

0

12

Warm and wet

H-events

1

18

1

13

17

10

Cold and dry

Bølling–Allerød

2

22

7

1

3

15

Warm and wet

Younger Dryas

1.3

35

3

12

14

17

Cold

The 8.2 ka event

0.16

13

0

16

6

6

Cold

MCA

0.4

28

11

0

24

10

Warm and dry

LIA

0.5

36

9

15

16

25

Cold and wet

out of five cold periods show a cold-and-dry regime, one
exhibits cold-and-wet conditions (Little Ice Age) and two
remain inconclusive (Younger Dryas and 8.2 k event).
It is easy to note that the periods that diverge from theClausius-Clapeyron thermodynamic response are theshorter ones (Fig. 10). Longer periods with durationcomparable to Holocene, such as the Eemian Interglacial
Stage and the Last Glacial Maximum follow the warmand-wet and cold-and-dry paradigms. A similar pattern
manifests in the spatial domain. Global or hemisphericstudies are more tightly linked to the thermodynamic
response, while as spatial scale becomes finer the heterogeneity increases highlighting the impact of the changes
in atmospheric and oceanic circulation (Gasse 2000; Liet al. 2012).

Thus, it is reasonable to claim that the atmospheric/oceanic circulation (dynamic component) appears tohave a more dominant role in the regional fluctuationsof the hydrological cycle, than the total atmosphericmoisture content (thermodynamic component). Thisis particularly true for the abrupt climatic events. Even
though the exact physical mechanisms of their genesis
remain under investigation, there is a general agreement that most of the past abrupt climatic transitions
are related to changes in the oceanic circulation. Still,
it is a concern whether these abrupt climate changes
arisen from internal climate system processes or be theconsequence of a stimulated response to a progressive
external forcing (Clement et al. 2001). In the case of
longer climatic regimes, warmer/colder oceans develop
different circulation patterns, which in turn affect theFig. 10 Schematic representation of hydroclimatic conditions in terms of the period length and the uncertainty involved. Uncertainty isqualitatively derived from the number of studies

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

atmosphere system resulting to different modes ofatmospheric circulation (Wright et al. 1992).
The majority of the abrupt events studied here weremainly associated with the Atlantic longterm variability
and AMOC in specific. The AMOC is not a circulation
pattern appearing only in Holocene. Its existence has
been confirmed both for the Eemian Interglacial Stageand the D–O intervals (Corrick et al. 2020). The declineof AMOC strength has also been linked to Heinrich
events, Younger Dryas, the 8.2 ka event, and phasesof cold conditions in general (Ellison et al. 2006; Renssen et al. 2018). It’s weakening is related to freshwater
pulses caused by the melting of Arctic ice and highlatitude glaciers (Li et al. 2012). The AMOC variability can affect the Westerlies, and, thus the atmosphericmoisture amount that is transferred over land. Whenweak, it has been linked to decline in precipitationfrom western Europe to continental Asia (Mackay et al.
2013), as well as monsoon activity (Gupta et al. 2003).
The latter is likely due to the links between the weakening of AMOC and the southward shift of Inter Tropical Convergence Zone (ITCZ). As the AMOC weakens,the temperature gradient between tropical and North
Atlantic becomes more intense and drives ITCZ to thesouth (Mohtadi et al. 2014).
Overall, the southward shift of the ITCZ has been
related to the colder conditions across the northern
tropics such as Heinrich events (Leduc et al. 2007),
Younger Dryas (Peterson and Haug 2006), and theweak monsoon during MIS-5e (McGowan et al. 2020).
In addition, the latitudinal variations of the ITCZ havebeen identified to affect summer-monsoon variations intropical and Asian regions during the D–O and Heinrich events (Ivanochko et al. 2005). On the contrary,
the northward shift of the ITCZ has been reported tointensify the Asian summer monsoon (Peterson et al.
2000; Wang et al. 2001), which is also related to warmer
conditions (Stansell et al. 2010; Schneider et al. 2020).
There is some evidence of this behaviour also during
mid-Miocene; the enhanced precipitation observed
across northern Colombia was likely due to the northward shift of ITCZ (Scholz et al. 2020). Most importantly, as ITCZ shifts the regions that are no longer
under its effect will become drier, with an opposite
outcome to the ones that no longer affected. This is astraightforward example of why wetter and drier conditions can co-exist when there is some atmosphericreorganization. As ITCZ and the monsoon systems
involve a large fraction of global precipitation, further
research is increasingly important to further understand the relationship between oceanic circulation and
ITCZ/monsoon in past climates.

Page 25 of 37

It is interesting that even though there is substantial
evidence of the connection between ITCZ and temperature in the paleoclimatic reconstructions, this was not thecase for atmospheric moisture divergence zones as well.
Nowadays, the dominant hypothesis suggests that global
warming makes the regions with atmospheric divergence
to become drier and the regions with atmospheric convergence to become wetter, termed as ’dry gets drier, wetgets wetter’ (Held and Soden 2006). However, our results
are not in favor of this hypothesis, which have also been
recently debated by some empirical studies of observational (Greve et al. 2014) and paleoclimatic records (Burls
and Fedorov 2017). On the other hand, we notice that inmany periods, the prevailing hydroclimatic regime, e.g.
warm and wet, appears in 65–80% of the studies. This
could imply that the convergence/divergence did become
stronger in the past warmer periods, but at the sametime a substantial reorganization of the atmospheric circulation patterns occurred. Plainly speaking the intensification did occur, but it might have affected different
regions.
To further investigate the spatial heterogeneity of thetemperature/precipitation relationship, we also examined the hemispheric and latitudinal distribution of therecords during cold and warm periods. No significant
changes are observed between the hemispheric distribution of studies during cold periods (Fig. 11A). In the zonal
domain, there is a divergence between mid and high latitudes, with the former exhibiting a tendency to cold-andwet conditions and the latter cold-anddry (Fig. 11B). In
addition, approximately one-third of the studies document warm climates over mid latitudes, with a higher
occurrence in SH. Even though the uneven number ofstudies per hemisphere and latitudinal zone makes theinterpretation of the results ambiguous, it provides someinsight of the enhanced heterogeneity, especially whencompared with the warm periods. The warm periods
appear quite more homogeneous in terms of temperaturefor both hemispheres (Fig. 12A). The NH appears to favorwarm-and-wet conditions in a 2:1 ratio, which drops toapproximately 1:1 for SH. The distribution appears quite
similar for all three latitudinal zones, also close to 2:1(Fig. 12B). Again, the bias of the low number of studies
in SH should be taken into account. Nevertheless, thehemispheric and latitudinal distribution of the records
advocate for an asymmetric response of precipitation totemperature increase and decrease. It should be noted
though that due to the limited number of records (especially precipitation), it is difficult to adequately describe
the spatial features of the water cycle’s response.
A plausible approach to overcome the reconstruction
scarcity barrier can be found in earth system modelling. Indeed, evidence of abrupt or mild atmosphericPratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 26 of 37

Fig. 11 Relationship of temperature and precipitation during the cold periods (LGM, Heinrich Events, Younger Dryas, 8.2 ka Event, and LIA). A
studies over different regions, B studies over the different zones

reorganization has been presented for some of the climatic periods discussed in our review. For instance,
model simulations show that during the MCA wasassociate with a substantial expansion of the NH Hadley circulation (Graham et al. 2011). This change theatmospheric circulation patterns could explain the drying over the mid-latitudes and the shifts in the monsoon patterns across Africa and South Asia. Other

reorganization patterns have been suggested for thetermination of the last deglaciation (Wassenburg et al.
2016) or the LGM (Justino et al. 2005). Even though
this evidence is far from conclusive, the hypothesis ofa circulation-modulated water cycle intensification is apromising direction to reconcile the ’dry gets drier, wetgets wetter’ paradigm with the observed changes.

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

Page 27 of 37

Fig. 12 Relationship of temperature and precipitation during the warm periods (MMCO, MIS5e, DO events, Bølling–Allerød interstadial, and MCA).
A studies over different regions, B studies over the different zones

Unfortunately, earth system modelling comes with certain limitations as well. The consistency between model
output and proxy data shows agreement over larger
scales, but there are crucial discrepancies in the regional
scales (Heiri et al. 2014). This is a known issue in model
performance related to the challenges in reproducing
precipitation properties at finer/regional scales (Flatoet al. 2014). However, some uncertainties still remain inlarger scales, due to inconsistencies in the simulation of
atmospheric circulation (Allan et al. 2020) and its modes
such as ENSO (Bellenger et al. 2014) or NAO (Zappa andShepherd 2017; Deser et al. 2017). Inevitably, the AMOC
is also poorly represented (Zhang et al. 2019), whichmight be related by a common bias in the model’s parameterization regarding AMOC stability (Liu et al. 2017).
On the other hand, the past millennium scale records

Pratap and Markonis ﻿Progress in Earth and Planetary Science(2022) 9:30

reveal no evidence of internally originated multidecadaloscillation. These multidecadal Atlantic Multi-decadal
Oscillation-like oscillations have contradicted as the
manifestation of high-amplitude explosive volcanism episodes (Mann et al. 2021). With all these open challenges
in earth system modelling, the need for more high-resolution paleoclimatic reconstructions is increasing.
More paleoclimatic reconstructions would further
improve our understanding the interaction between temperature and water cycle. The evidence presented heresuggest that the hypothesis that a warmer climate is awetter climate could be an oversimplification even for
centennial scales. On the contrary, precipitation response
appear to be spatio-temporally heterogeneous, with certain differences among periods. This should be taken into
account when assessing the future intensification of the
global water cycle. Even if not regionally precise, the precipitation response heterogeneity should be evident inmodel simulations or our theoretical constructs of the
global water cycle functioning. This qualitative metric
could help improve the model performance, and in turnshed more light on the influence of atmospheric and oceanic circulation. The remaining challenge, though, is toquantify the spatial variability of precipitation response
in a robust manner. As the number of paleoclimaticreconstructions increases, we will soon be able to have amore coherent picture of specific warm or cold periods,
and increase the likelihood to address it.

6 Conclusions
Most climate projections report that the hydrologicalcycle will intensify when the climate will get warmer. As
a result, the hydrological cycle sensitivity is a major concern for the coming decades. In this study, we reviewed
the relationship between the hydroclimate and temperature in the recent and distant past. We confirmed that,in general, most paleoclimate records suggest that thehydrological cycle intensified in a warmer climate. Correspondingly, the hydrological cycle weakened during
the colder periods. However, the spatial distribution ofhydroclimatic changes was not homogeneous around theworld.
This lack of homogeneity makes paradigms such as “awarmer climate is a wetter climate” or “dry gets drier,
wet gets wetter” appearing as oversimplifications. The
evidence presented in this study agrees to the hypothesisthat climate changes at global scale are thermodynamicdriven, while regional climate changes are more related
to variations in ocean-atmospheric circulation. However, due to its enhanced spatiotemporal distribution,hydroclimate variability is difficult to be quantified ona regional, continental, and global scale. In this context,
large-scale paleo-hydroclimatic shifts, especially during

Page 28 of 37

the warm periods, need further investigation as they
could provide new insights into the present and futurehydroclimatic changes.
Acknowledgements
The authors would like to thank the two anonymous reviewers for theirconstructive remarks that helped to improve the manuscript. This work was
carried out within the project “Investigation of Terrestrial HydrologicAl Cycle
Acceleration (ITHACA)” funded by the Czech Science Foundation (Grant
22-33266M). SP was supported by the Internal Grant Agency of the Facultyof Environmental Sciences, Czech University of Life Sciences Prague (Grant
2020B0039).
Author contributions
SP and YM conceived the idea and designed the study. SP carried out theliterature research, analyzed the data, and prepared the manuscript withsupport from YM. Both authors discussed the results and approved the final
manuscript.
Funding
This study was financially supported by the Czech Science Foundation (Grant
22-33266M) and the Internal Grant Agency of the Faculty of Environmental
Sciences, Czech University of Life Sciences Prague (Grant 2020B0039).
Availability of data and materials
The datasets used for this study are obtained from NOAA’s National Centers
for Environmental Information webdata portal, available at https://​www.​ncei.​noaa.​gov/​produ​cts/​paleo​clima​tology. Another source for paleoclimate data is
Past Global Changes (PAGES) databases, available at https://​www.​pastg​lobal​
chang​es.​org/​scien​ce/​data/​datab​ases.

Declarations
Competing interests
The authors declare that they have no competing interest.
Received: 6 December 2021 Accepted: 16 May 2022

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