M E D D E L A N D E N

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S T O C K H O L M U N I V E R S I T E T S I N S T I T U T I O N

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G E O L O G I S K A V E T E N S K A P E R

N o . 3 5 3

Asian monsoon over mainland Southeast Asia

in the past 25 000 years

Akkaneewut Chabangborn

Stockholm, October 2014

Department of Geological Sciences Stockholm University SE-106 91 Stockholm

Sweden

©Akkaneewut Chabangborn, Stockholm 2014 ISBN 978-91-7447-969-0 Printed in Sweden by US-AB Stockholm University, Stockholm, 2014 Cover: Naga bridge from Prasat Phanom Rung, Buriram, Thailand. The Naga is a symbol for rains and water.

Asian monsoon over mainland Southeast Asia

in the past 25 000 years

Akkaneewut Chabangborn

Department of Geological Sciences, Stockholm University, SE-106 91, Stockholm, Sweden

This thesis consists of a summary, and three published articles (Paper I-III) and one submit-ted manuscript (Paper IV);

Paper I

Chabangborn, A., Brandefelt, J., Wohlfarth, B., 2013. Asian monsoon climate during the Last Glacial Maximum: palaeo-data–model comparisons. Boreas 43, 220-242.

Paprt II

Chabangborn, A., Wohlfarth, B., 2014. Climate over mainland Southeast Asia 10.5–5 ka. Jour-nal of Quaternary Science 29, 445-454.

Paper III

Chawchai, S., Chabangborn, A., Kylander, M., Löwemark, L., Mörth, C.M., Blaauw, M., Klubse-ang, W., Reimer, P.J., Fritz, S.C., Wohlfarth, B., 2013. Lake Kumphawapi - an archive of Holo-cene palaeoenvirontal and palaeoclimatic changes in northeast Thailand. Quaternary Science Reviews 68, 59-75.

Paper IV

Chawchai, S., Chabangborn, A., Fritz, S., Väliranta, M., Mörth, C.M., Blaauw, M., Reimer, P.J., Krusic, P.J., Löwemark, L., Wohlfarth, B., Hydroclimatic shifts in northeast Thailand during the last two millennia – the record of Lake Pa Kho. Submitted to Quaternary Science Reviews.

To my parents and my wife

A dissertation for the Degree of Doctor of philosophy in Natural Science

Akkaneewut Chabangborn Department of Geological Sciences Stockholm University, SE-106 91, Stockholm, Sweden

Abstract Mainland Southeast Asia is located close to the boundary between the Indian Ocean and

the East Asian monsoon sub-systems. Strong convective cells develop over this region during the initial stage of the summer monsoon onset; the start of the summer monsoon season is thus earlier than in neighboring areas. The main objective of this research study is to contrib-ute, analyse, and interpret high-resolution palaeo-proxy data sets from this region to further understanding of Asian summer monsoon variability in the past. This was done by compiling, evaluating, and synthesizing published palaeo-records from the Asian monsoon region; through data – model simulation comparisons; and by analysing new lake sedimentary re-cords from northeast Thailand.

Palaeo-records and climate model output indicate a strengthened summer monsoon over Mainland Southeast Asia during the Last Glacial Maximum at 23 000-19 000 years before present (BP), as compared to dry conditions in other parts of the Asian monsoon region. This can be explained by the Last Glacial Maximum sea level low stand, which exposed the conti-nental shelf of Sundaland and created a large land-sea thermal contrast. The rapid sea level rise ~19 600 years BP led to the reorganization of the atmospheric circulation in the Pacific Ocean and possibly played an important role in the weakening of the summer monsoon be-tween 20 000 and 19 000 years BP.

Reconstructions of semi-quantitative precipitation and temperature for Mainland South-east Asia show that annual rainfall amounted to more than 1100 mm and that annual tem-peratures were above 18°C between 15 000 and 5000 years BP. The hydroclimatic records for Mainland Southeast Asia, together with those established for the East Asian monsoon sub-region, moreover point to an earlier Holocene onset of a strengthened summer monsoon, as compared to the Indian Ocean monsoon sub-region. The asynchronous evolution of the sum-mer monsoon and the time lag of about 1500 years between the East Asian and the Indian Ocean monsoon sub-regions can be explained by the palaeogeography of Mainland Southeast Asia, which acted as a land bridge for the movement of the Intertropical Convergence Zone.

The new palaeo-proxy records from northeast Thailand compare well to the semi-quantitative data sets and suggest a strengthened summer monsoon between 10 000 and 7000 years BP and a weakening of the summer monsoon thereafter. Warmer temperatures in the Northern Hemisphere during the early Holocene led to a northward shift in the mean po-sition of the Intertropical Convergence Zone, while the mean position of the Intertropical Convergence Zone moved southward during the mid Holocene following the decrease in inso-lation. The multi-proxy, high-resolution lake/peat sequence from Lake Pa Kho provides a picture of summer monsoon variability during the last 2000 years. A strengthened summer monsoon prevailed between BC 170 and AD 370, between AD 800 and 960 and since AD 1450, while it was weaker about AD 370-800 and AD 1300-1450. Although these shifts in summer monsoon intensity can generally also be explained by the movement of the mean position of the Inter-tropical Convergence Zone, the weakening of the summer monsoon between AD 960 and 1450 was possibly affected by changes in the Walker circulation.

A dissertation for the Degree of Doctor of philosophy in Natural Science

Akkaneewut Chabangborn Department of Geological Sciences Stockholm University, SE-106 91, Stockholm, Sweden

Sammanfattning

Sydostasien är beläget där den Indiska Oceanen monsunen möter den Öst-Asiatiska monsunen. Därför bildas starka konvektionsceller över denna region i början av sommarmonsunen, som inleds i ett tidigare skede här jämfört med i angränsande områden. Målet med denna vetenskapliga studie är att bidra, analysera och tolka paleo-proxy data från denna region för att få en fördjupad förståelse av hur den Asiatiska sommarmonsunen har utvecklats fram till idag. Detta mål nåddes genom att sammanställa och utvärdera redan pulblicerade paleo-data från den Asiatiska monsunregionen; men också genom att jämföra data-modell simulationer och genom att analysera nya paleo-proxies från sjösediment från nordöstra Thailand. Resultat från paleo-data och klimatmodeller indikerar en förstärkt sommarmonsun över Sydostasien under senaste istidsmaximum mellan 23 000 och 19 000 år före vår tideräkning (BP) (innan AD 1950), jämfört med torrare förhållanden i andra delar av den asiatiska monsunregionen. Denna skillnad kan ha sin förklaring i havsnivån, som var betydligt lägre under senaste istidsmaximum, vilket blottade kontinentalsockeln Sundaland och skapade en termalt kontrast mellan land och hav. Den snabba stigningen av havsnivån ~19 600 år BP ledde till en omorganisering av den atmosfäriska circulationen i Stilla havet, vilket kan ha bidragit till försvagningen av sommarmonsunn mellan 20 000 och 19000 år BP. Semikvantitativa rekonstruktioner av variationer i nederbörd och temperatur för Sydostasien visar att den årliga nederbörden var mer än 1100 mm och att temperaturen var över 18°C mellan 15 000 och 5000 år BP. I övrigt indikerar hydroklimatiska proxys för Sydostasien tillsammans med data från den östasiatiska monsunregionen att den förstärkta sommarmonsunen under Holocen påbörjades tidigare än monsunen ovanför den Indiska Oceanen monsunregionen. Den asynkrona utveckligen av sommarmonunen och tidsskillnaden på 1500 år mellan den Östasien och Indiska Oceanen monsunregionen kan förklaras av palaeogeografin av sydostasien, som fungerade som en landbrygga för den intertropiska konvergenszonen. De nya plaeo-proxy data från nordöstra Thailand är välkorrelerade med de tidigare nämnda semikvantitativa dataseten och föreslår en förstärkt sommarmonsun mellan 10 000 och 7000 år BP samt en försvagning av den efterföljande sommarmonsunen. Varmare temperaturer i den norra hemisfären under tidig Holocen ledde till en nordlig skiftning i medelpositionen för den intertropiska konvergenszonen. Den minskade solinstrålningen under mitten av Holocen gjorde att medelpositionen för den intertropiska konvergenszonen flyttade söderut under denna tid. Multiproxystudien från sjö/torvsekvensen i Pa Kho sjön ger en indikation av variabiliteten av sommarmonsunen under de senaste 2000 åren. En förstärkt sommarmonsun pågick från BC 170 och AD 370, mellan AD 800 och 960 och efter AD 1450, medans den var svagare mellan AD 370-800 och AD 1300-1450. Trots att dessa intensitetsförändringar i sommarmonsunen generellt sett kan förklaras genom förflyttnignen av medelpositionen av den intertropiska konvergenszonen så kan försvagningen av sommarmonsunen mellan 960 och 1450 möjligen bero på förändringar i Walkercirkulationen.

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Contents 1 Introduction 9 1.1 The Asian monsoon – a short overview 9 1.2 The Asian monsoon during the late Quaternary and

outstanding questions

12 1.3 Environmental setting in MSEA 12 1.4 Land-sea configuration during the late Quaternary 14 1.5 Palaeoclimate studies in mainland Southeast Asia 16 1.6 Thesis objectives 16 2 Methods of study 18 2.1 Palaeo-records compilation 18 2.1.1 Palaeo-proxy assessment in Paper I 18 2.1.2 Reconstruction of semi-quantitative precipita-

tion and temperature in Paper II

21 2.2 Lake sediment analysis 23 3 Summary of the papers 25 4 Results 26 4.1 Paper I: Asian monsoon climate during the Last Gla-

cial Maximum: palaeo-datamodel comparisons

25 4.2 Paper II: Climate over mainland Southeast Asia 10.5-5

ka BP

29

4.3 Paper III: Lake Kumphawapi – an archive of Holocene palaeoenvironmental and palaeoclimatic changes in northeast Thailand

32 4.4 Paper IV: Hydroclimate shifts in northeast Thailand

during the last two millennia – the record of Lake Pa Kho

34 5 Discussion 36 5.1 The last glacial maximum (23-19 ka BP): Paper I 36 5.2 The early to mid Holocene (10.5-5 ka BP): Paper II

and III

37 5.3 The late Holocene (the last 2 ka): Paper IV 40 6 Conclusions 43 7 Ongoing works and future perspective 44

7.1 Asian monsoon variability during the marine isotope stage 3 (MIS 3)

44

7.2 Myanmar Project 45

8 Acknowledgements 47 9 References 48 Appendix I i Appendix II iii

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1. Introduction

1.1 The Asian monsoon – a short overview The monsoon is considered a primary component of the global climate system as it transfers

heat energy and humidity from the equator to higher latitudes (e.g. Webster et al., 1998; An et

al., 2000; Trenberth et al., 2000; Zhan, 2003; Cook and Jones, 2012; Jo et al., 2014). It is gener-ally described as a seasonal reversing wind associated with distinct changes in precipitation, which is caused by land-sea thermal contrasts (e.g. Ramage, 1971; Wang et al., 2003; Wang et

al., 2005; Wang, 2009; Clift and Plumb, 2008). Monsoon variability is fundamentally controlled by changes in insolation (external forcing) and by land, ocean and atmosphere interactions (internal forcing) (Clemens et al., 1991). In summer, the increase in insolation leads to the de-velopment of strong convective cells over the continent, which allows high humidity air masses to flow from the ocean to the land. The incursion of humid air masses enhances the amount of precipitation over the monsoon region. A decrease in insolation and relatively warm sea sur-face temperatures (SST) in winter cause a reversal of the winds from the continent to the oceans and leads to cold and dry climate conditions over the monsoon region.

The different monsoon regions have been generally classified according to seasonal differ-ences in surface wind components and are confined to South and East Asia, North Africa, and Australia (Ramage, 1971; Black, 2002; Wang et al., 2005; Wang, 2009; Wang et al., 2014) (Fig. 1A). Recent studies, which are based on the development in meteorological observations and satellite technology, however introduce a new concept, the ‘global monsoon’ (Trenberth et al., 2000; Wang, 2009; Wang et al., 2014); following these studies, monsoon regions are defined by seasonal changes in precipitation associated with the migration of the rain belt following the monsoon trough and the Intertropical Convergence Zone (ITCZ). Since the ITCZ is a global-scale, low-latitude feature, the ‘global monsoon’ concept (Wang, 2009) includes six monsoon regions: Asia, Australia-Indonesia, North and South Africa, and North and South America (Figs. 1B, 1C). Among these six monsoon regions, the Asian and African monsoons are regarded as the most prominent and interesting (Wang, 2009).

Here I follow Wang et al. (2003) and Wang et al. (2005) to constrain the Asian monsoon re-gion; it extends from the Arabian Sea to the South China Sea and from northern Australia to northern China (Fig. 2). According to this definition, the Asian monsoon controls the annual hydroclimate variability over an area, which is a home to more than half of the world’s popula-tion (An et al., 2000; Wang et al., 2003; Wang et al., 2005; Clift and Plumb, 2008; Buckley et al., 2014). Most of the people living in this part of the world work in agriculture and their produce is strongly dependent on the intensity of annual rainfall (Webster et al., 1998; An et al., 2000; Clift and Plumb, 2008; Buckley et al., 2014). Moreover, the Asian monsoon is the largest of the different monsoon systems; its significant northward incursion onto the continent (~40°N) makes it a potentially important player linking low and high latitudes (Fig. 1A). Rainfall vari-ability over a region with a high population density and the importance of the monsoon as a global-scale climate system makes the monsoon highly sensitive to future climate change; negative or positive changes will have large regional and global scale effects on environment, society and economy (IPCC 2012).

The Asian monsoon system is generally divided into the Indian Ocean (IOM) and East Asian monsoon (EAM) sub-systems, according to summer wind circulation patterns (Wang et al., 2003; Wang et al., 2005) (Fig. 2). The IOM is described by a distinct clockwise circulation across the equatorial Indian Ocean and associated with the Somali Jet and southeasterly winds generated by the Mascarene High. The EAM is sometimes divided into the EAM and the West-ern North Pacific Monsoon (WNPM), because of the significant relationship between the west-ern Pacific Warm Pool and the El Niño Southern Oscillation (ENSO) (Wang et al., 2003). Since changes in ENSO do not occur on an annual basis and since it originates in the eastern Pacific and outside of the Asian Monsoon region, I here consider this factor as a remote forcing. Con-sequently, the EAM sub-system can be regarded as a convergence of the Pacific trade wind, the

10

subtropical front near the east coast of China and the southwest monsoon from the Indian Ocean. The boundary of the IOM and EAM sub-systems extends along longitude 105°E, i.e. from the eastern edge of the Tibetan Plateau through mainland Southeast Asia and into the Indone-sian Archipelago (Fig. 2).

Fig. 1: (A) Distribution of modern monsoon regions (orange shaded) based on the seasonal reversal of winds (Ramage, 1971; Wang et al., 2005); (B) mean position of the Intertropical Convergence Zone (ITCZ) (green lines) in summer and winter. (C) The different monsoon re-gions according to the Global Monsoon Concept (Wang, 2009): (a) North America, (b) South America, (c) North Africa, (d) South Africa, (e) Asia and (f) Australia-Indonesia (yellow shaded).

Modern observations show that the Asian monsoon is a much more complicated system than a simple large-scale land-sea breeze, as it is modulated by various factors, e.g. ENSO and/or the Western Pacific Warm Pool (Goswami and Xavier, 2005); the Indian Ocean Dipole (IOD) (Weller and Cai, 2014); the Atlantic Meridional Overturning Circulation (AMOC) (Yu et

al., 2009); the movement of the ITCZ (Chao and Chen, 2001); change in SST (Li et al., 2001) and atmospheric constituents (Ramanathan et al., 2005).

Climate models simulate an increase in land-sea thermal contrast and a subsequent strengthening of the summer monsoon under future global warming scenarios (IPCC 2012). However, IPCC (2012) stressed that there is low confidence regarding future changes in the Asian monsoon because the different climate models provide highly divergent output. Mete-orological observations and data analysis (1950-2000) however provide a more differentiated

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picture. Annual precipitation has increased over the western part of China and the southeast coast of China, over the northeastern Tibetan Plateau and the Philippines, and decreased over Southeast Asia; the patterns over India are uncertain (Ramahathan et al., 2005; Cruz et al., 2007; Zhang et al., 2011).

Fig. 2: Association of mean precipitable water content (in colour) and wind circulation pat-terns (arrows) in winter (January and February) and in summer (June, July and August) based on the reanalysis of the National Centers for Environmental Prediction and the National Center for Atmospheric Research (NCEP/NCAR) from 1981 to 2010. The Indian and East Asian mon-soon sub-regions are separated by the white dashed line (105°E).

These uncertainties and discrepancies necessitate a better understanding of the long-term

hydroclimate history in the region, which allows placing modern climate change into a geologi-cal and historical perspective. Extreme climate states in the past (e.g., the Last Glacial Maxi-mum; Heinrich events), which are recorded in geological archives, moreover provide a possi-bility to test the performance of climate model and the influence of external and internal forc-ing.

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1.2 The Asian monsoon during the late Quaternary and outstanding ques-

tions Palaeo-studies are mainly restricted to the late Quaternary because this time interval offers

more reliable geochronologies and has also been relatively stable from a tectonic perspective. Reconstructed cool and dry conditions between 25 and 17 thousand years (ka) before present (BP) have conventionally been interpreted as a weakening of the Asian summer monsoon (e.g. An et al., 1991; Chen et al., 1997; Huang et al., 1997). Some palaeo-records moreover suggest that cool and dry conditions prevailed over the Asian monsoon region until 12 ka BP (e.g. Bhattacharyya et al., 1989; van der Kaars, 1997; Huang et al., 1999; Anshari et al., 2001; van der Kaars et al., 2001; Mingram et al., 2004). Other studies however point to a gradual increase in precipitation and a stronger summer monsoon between 17 and 12 ka BP (e.g. Vasathy et al., 1988; An et al., 1991; Yang et al., 1995; Chen et al., 1997; Rajagopalan et al., 1997; Williams et

al., 2008). Most palaeo-records agree regarding an abrupt increase in precipitation during the early Holocene due to a strengthening of the summer monsoon (e.g. Morinaga et al., 1993; Chen et al., 1999; Hodell et al., 1999; Huang et al., 1999; Dam, 2001; van der Kaars et al., 2001; Wang et al., 2002; Mingram et al., 2004; Zhang et al., 2009). After 7 ka BP, the Asian summer monsoon intensity started to weaken due to the overall decrease in insolation (Yang et al., 1995; Bernasconi et al., 1997; Yan et al., 1999; Mingram et al., 2004; Wang et al., 2010a; Ste-venson et al., 2011).

Recent studies with better chronologies however show that the general trend of a strength-ening/weakening of the Asian monsoon was punctuated by several shorter intervals of abrupt climate shifts. Short intervals of wetter conditions are for example reported for 25 to 17 ka BP (e.g. Rashid et al., 2007; Saher et al., 2007; Govil and Naidu, 2011; Mahesh et al., 2011) and 14 to 13 ka BP, and drier intervals between 12.5 and 11.5 ka BP and ~8 ka BP (Dykoski et al., 2005; Sinha et al., 2005; Shakun et al., 2007). Palaeo-records from China moreover discuss an asynchroneity between the IOM and EAM sub-systems in respect to the onset of the summer monsoon during the early Holocene (An et al., 2000; Herzschuh, 2006; Maher, 2008; Wang et

al., 2010b). These new studies highlight the fact that the temporal variability of the Asian mon-soon is still an open question.

Several of these palaeo-records have been assembled and synthesized on a regional scale, e.g. China (Herzschuh, 2006; Wang et al., 2010b), India (Tiwari et al., 2003 and 2011), South-east Asia (Cook and Jones, 2012), Indonesia, and the northern part of Australia (Reeves et al., 2013). Since the Asian monsoon affects a vast area and since each of these compilation studies used different techniques, proxies and assumptions, valid spatial comparisons across the en-tire Asian monsoon region are limited.

Although summer monsoon intensity is thought to follow changes in insolation, most of the palaeo-studies agree that the peak in summer monsoon intensity lags maximum insolation by approximately 3 ka (Overpeck et al., 1996). The reason for this lag between maximum summer monsoon intensity and insolation has been debated and has been attributed to a number of factors. ENSO, atmosphere-ocean interactions (Wang et al., 2005), a southward shift of the ITCZ (Broccoli et al., 2006; Zhang and Delworth, 2005; Braconnot et al., 2007), the Pacific Warm Pool (De Deckker et al., 2002; Partin et al., 2007), and North Atlantic circulation (Over-peck et al., 1996; Tiwari et al., 2009; Pausata et al., 2011; Stager et al., 2011) may have played an important role in modulating Asian monsoon variability through time.

Several of these outstanding questions are likely to be related to a lack of high-resolution palaeo-proxy studies from mainland Southeast Asia, which is considered as a key location for the study of the Asian monsoon (e.g. Cook and Jones, 2012; Buckley et al., 2014).

1.3 Environmental setting in MSEA

Southeast Asia is conveniently divided into mainland (continental) and maritime Southeast Asia based on its geographical setting. Mainland Southeast Asia (MSEA) is characterised by

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more seasonal climate variability than over the maritime continent of Southeast Asia, which experiences more uniform humid conditions (Chuan, 2005). MSEA covers the area of present day Myanmar, Thailand, Laos, Cambodia, Vietnam, the Malay Peninsula and Singapore.

Orogeny has played a major role in shaping the topography of MSEA. The series of collisions between the Indian and Eurasia plates since the Eocene, accompanied by climate change dur-ing the Quaternary, are the main factors behind the geographical development of MSEA (Gupta, 2005). The northern and northwestern part of MSEA is made up of high mountain areas, pla-teaus and hills, which form part of the Himalayas and which are intersected by deep valleys (Gupta 2005). The central mountain ranges extend into Myanmar and Thailand, along the Ma-lay Peninsula, and into the Indonesian archipelago, while the eastern range continues from northern Laos to southern Vietnam. The major rivers, Irrawaddy, Chao Phraya, Mekong and Sông Hóng (Red River), created wide alluvial plains between the mountain ranges. An excep-tion is the region between the Chao Phraya Plain and the Mekong Lowland, which is separated by the gently sloping Khorat Plateau. The major MSEA rivers originate on the eastern Tibetan plateau, except for the Chao Phraya River, which has its source in the MSEA region.

Fig. 3: (A) Mean position of the Intertropical Convergence Zone (ITCZ) and path of tropical cy-clones (dashed orange line) over MSEA for each month. The red and blue arrows mark the southwest and northeast mon-soon, respectively. The dashed black line (longitude 105°E) marks the boundary between the Indian Ocean (IOM) and East Asian (EAM) monsoon sub-systems. (B) Monthly rainfall data for 1981 to 2010 shows a distinct difference in seasonal precipitation over Thailand (data from Thai Meteorological Department) (B).

14

Based on the Köppen-Geiger climate classification, MSEA is assigned to a tropical climate system that is characterised by high annual precipitation, and annual temperatures exceeding 18°C (Peel et al., 2007). Seasonal temperature difference is very small in most of the MSEA region but increases with latitude (Chuan 2005). The National Center for Atmospheric Re-search (NCAR) Community Climate Model (CCM1) simulates mean annual temperatures of ~25°C, precipitation of 3170 mm a-1 and effective moisture (precipitation minus evaporation) of 1825 mm a-1 over MSEA and the maritime continent (Zhang et al., 1996). Annual precipita-tion over MSEA is mainly controlled by the Asian monsoon, the seasonal position of the ITCZ and receives contributions from tropical cyclones during the summer/winter and win-ter/summer transition (Takahashi, 2011; Reid et al., 2013). The establishment of strong con-vection and reversing winds over the Bay of Bengal, MSEA and the South China Sea occurs dur-ing the Asian summer monsoon onset in May (Matsumoto, 1997; Zhang et al., 2005; Buckley et

al., 2014). These allow a northward shift of the ITCZ, winds blowing from the surrounding oceans to the continent and transport of humidity (Fig. 3A). The ITCZ reaches the northern part of MSEA in September, and moves southward about a month later. This leads to high pre-cipitation over MSEA between May and October (Fig. 3B). The precipitation amount induced by the Asian summer monsoon, however, also receives a strong influence from tropical cyclones that develop over the west Pacific Ocean, the South China Sea and the north Indian Ocean dur-ing the transition from summer to winter and vice versa (Takahashi, 2011). A southward shift of the ITCZ allows the penetration of cold and dry air masses from Siberia, which leads to rela-tively drier conditions in MSEA between November and February (Figs. 3A, 3B).

1.4 Land-sea configuration during the late Quaternary Dramatic sea level changes during the late Quaternary strongly altered the land-sea configu-

ration in MSEA (Fig. 4). Since the land-sea distribution is a primary driver of monsoon variabil-ity, significant changes in MSEA’s palaeogeography may have had an important impact on land-ocean-atmosphere interactions and past Asian monsoon dynamics.

During the last glacial maximum (LGM) (23-19 ka BP) the sea level attained a low stand of ~125 m below present levels (BPL), but increased on average by about 10 m/ka between 19 and 8 ka BP (Voris, 2000; De Deckker et al., 2002; Bird et al., 2005; Clack et al., 2009; Hanebuth et al., 2011) (Fig. 4A). However, this averaged sea level rise was modified by many short inter-vals of rapid increase in sea level, e.g. at 19.6 ka BP (15 m); 14.5 ka BP (20 m); 8.5 ka BP (5 m); and 7.5 ka BP (5 m) (Hanebuth et al., 2011) (Fig. 4A). At ~4.5 ka BP sea level reached its high-est stand of ~5 m above present levels and subsequently decreased toward the present level prior to 2 ka BP (Hanebuth et al., 2011) (Fig. 4A).

Advances in topographic and bathymetric observations allow generating approximate shoreline configurations for the late Quaternary (Figs. 4B-4G). The sea level low stand of ~125 m BPL at ~23 ka BP exposed the Sunda shelf and connected MSEA to Borneo, Sumatra, and Java (Voris, 2000; De Deckker et al., 2002; Bird et al., 2005; Hanebuth et al., 2011) (Fig. 4B). This emerged area was potentially characterised by a low gradient and three major river sys-tems: (i) the Malacca Straits river between the Malay Peninsula and Sumatra running in northwest-southeast direction to the Andaman Sea; (ii) the Siam river extending from the Chao Phaya lowland across the Gulf of Thailand in northwest-southeast direction; and (iii) the North Sunda river flowing from Sumatra to the Gulf of Thailand in northeast-southwest direction (Voris, 2000; Sathiamurthy and Voris, 2006) (Fig. 4B).

15

Fig.

4: (

A)

Sea

leve

l cu

rve

for

MSE

A f

or t

he

pas

t 2

1 ka

BP

(H

aneb

uth

et

al.,

20

11

); (

B)

pal

aeo-

shor

elin

e at

20

ka

BP

for

a s

ea le

vel l

ow s

tan

d o

f 1

25

m b

elow

pre

sen

t le

vel (

BP

L), a

nd

at

(C)

14

.3, (

D)

12

.7, (

E)

10

.8 k

a B

P, (

F) 9

.5, a

nd

(G)

4.5

ka

BP

for

sea

leve

l of

80

, 4

0, 2

0 m

BP

L, a

nd

5 m

ab

ove

pre

sen

t le

vel,

resp

ecti

vely

(m

odif

ied

from

Vor

is, 2

00

0 a

nd

Han

ebu

th e

t a

l., 2

01

1).

16

The palaeo-shoreline at ~14.3 ka BP was at 80 m BPL and was fairly similar to the sea level low stand at 23 ka BP. However, a significant transgression occurred near the Siam and North Sunda estuaries (Fig. 4C). This transgression migrated further inland near the Siam, North Sunda and Malacca Straits estuaries ~12.7 ka BP (Fig. 4D). A sediment core obtained off the east coast of the Malay Peninsula indicated that the freshwater lakes or swamps possibly ex-isted in the Gulf of Thailand (Emery and Nino, 1963; Biswas, 1973; Voris, 2000). A significant change occurred at 10.8 ka BP when sea level was about 40 m BPL (Fig. 4E). Although this transgression significantly incised into the Gulf of Thailand, MSEA was still connected to Suma-tra and Java. The connection between Borneo and Sumatra became narrower, or a long narrow channel possibly separated them (Voris, 2000). MSEA became likely separated from Sumatra, Java and Borneo at 9.5 ka BP when sea level reached about 20 m BPL (Fig. 4F). The palaeo-shoreline is slightly difference from present day during the sea level high stand (5 m above present levels) at ~4.5 ka BP (Fig. 4G). This transgression was only dominant near the main river deltas, i.e. the Irrawaddy, Chao Phraya, and Mekhong Rivers. 1.5 Palaeoclimate studies in mainland Southeast Asia

Mainland Southeast Asia (MSEA) is an important area for Asian monsoon studies because it is located close to the boundary between the IOM and EAM sub-systems and is the only conti-nental area that develops strong convective cells during the initial state of the summer mon-soon onset (Zhang et al., 2004; Buckley et al., 2014). However, only a few low-resolution pa-laeo-records are available for MSEA (see for example Kealhofer and Penny, 1998; Penny 2001; White et al., 2004; Penny and Kealhofer, 2005; Cook et al., 2010; Wohlfarth et al., 2012 for lit-erature overviews). In addition, these studies focused mainly on pollen, coastal change or ar-chaeology, and were published in local sources only (e.g. Pramojanee and Hastings, 1983; Chaimanee, 1989; Udomchoke, 1989; Maloney and McAlister, 1990).

The palaeo-vegetation studies from Thailand by Kealhofer and Penny (1998), Maloney (1999), and Penny (2001), from Cambodia by Maxwell (2001) and Penny (2006), and from Vietnam by Proske et al. (2011) provide an overview of past environmental change based on the succession and distribution of vegetation. Cook and Jones (2012) compiled these studies and assessed them in terms of semi-quantitative climate change during the last 30 ka. Their reconstruction suggests cool/dry and moderate cool/dry conditions from 30 to 13 ka BP, mod-erate warm/wet and warm/wet conditions from 13 to 6 ka BP, and moderate warm/wet and moderate cool/dry conditions from 6 to 0 ka BP (Cook and Jones, 2012). These palaeoenvi-ronmental reconstructions however assume continuous deposition based on the few poor chronologies and thus dismiss breaks in deposition, which are frequently found in recent stud-ies (e.g. Smittenberg et al., 2011; Wohlfarth et al., 2012; Chawchai et al., 2013). Moreover, the general assumption of cold/dry or warm/wet conditions is possibly an overestimation since palaeo-proxies may respond to either temperature or precipitation. 1.6 Thesis objectives

This thesis research is part of the Asian monsoon project (Table 1). The specific focus of my thesis is:

� To contribute new well-dated lake/peat sequences and high-resolution proxies in MSEA

� To better understand hydroclimate patterns reflected in Late Quaternary Asian summer monsoon variability.

My contribution to the Asian monsoon project is detailed in Table 1.

17

Table 1: Summary of my contribution to the Asian monsoon project. The location of the differ-ent sites is shown in Figure 5. 1Included in this thesis; 2ongoing project; 3no sediments present in lake basin; 4future studies.

Site of studies Fieldwork

and/or coring Laboratory

work Data interpre-

tation

Figures and Ta-

bles 1. Lake Kumphawapi1,2 x x x x 2. Lake Pa Kho1,2 x x x x 3. Lake Han Sakon Nakhon3 x 4. Lake Laung3 x 5. Lake Kway3 x 6. Lake Phayao3 x 7. Lake Leng Sai2 x x x 8. Lake Sifai3 x 9. Lake Boraphet3 x 10. Sam Roy Yod2 x 11. Lake Thale Pron2 x x x 12. Lake Thale Song Hong3 x 13. Lake Thale Bun3 x 14. Cave exploration for speleothems2

x

15. Myanmar project4 x x

Fig. 5: Location of the sites investigated within the Asian monsoon project in MSEA. The red squares represent the palaeo-records included in this research thesis and the yellow squares mark ongoing projects. The blue circle shows the location of a future project in northwest Myan-mar. The lake sites with no sediment present are marked with white circles. See Table 1 for details.

18

2. Methods of study The thesis consists of three published articles and one submitted manuscript. Papers I and II are compilations of palaeo-records from the entire Asian monsoon region during the Last Gla-cial Maximum (LGM) (23-19 ka BP) and from MSEA between 10.5 and 5 ka BP. Papers III and IV focus on a reconstruction of hydroclimate conditions using multi-proxy lake sediments from Thailand, which is located almost in the middle of MSEA. 2.1 Palaeo-records compilation Published palaeo-records were assembled based on the criteria detailed in Table 2. All pub-lished 14C dates were calibrated with the Calib 6.0 online program (http://calib.qub.ac.uk/calib/calib.html) (Reimer et al., 2009). Age-depth curves were further created for each of the sites listed in Paper I. However, for Paper II, the poor chronology of sev-eral of the MSEA palaeo-records did not allow creating valid age-depth curves. We therefore only used pollen-stratigraphic intervals that are covered by a 14C date. Table 2: Selection criteria for the palaeo-data set compilation of Papers I and II.

Paper I Paper II

1. Published marine and terrestrial proxy records from the Asian monsoon region (15°S-40°N, and 40−160°E) (see Fig. 6) covering the Last Glacial Maximum were selected (see details in Appendix I).

1. Published Holocene pollen-stratigraphic records from Thailand, Cam-bodia, Vietnam, the Malaysian Peninsular, and Singapore were selected (see Fig. 7 and details in Appendix II).

2. Each record should contain at least one 14C, U/Th, and/or TL date between 23 and 19 ka BP.

2. Each record should contain at least one 14C age estimate between 10.5 and 5 ka BP.

3. Records with age errors of >1000 years and 14C dates on carbonate bulk sediment were excluded.

3. The following pollen taxa were ex-cluded: mangrove and aquatic; Pinus and Poaceae; and those not included in the Southeast Asian and Pacific (SEAPAC) data set (Pickett et al., 2004).

2.1.1 Palaeo-proxy assessment in Paper I

The various palaeo-proxies (pollen, faunal remains, δ18O of speleothems and planktonic foraminifera, δ13C measured on bulk sediment) were assessed in terms of qualitative effective moisture (P-E) and precipitation during the LGM (23-19 ka BP) (Clark et al., 2009). The ob-tained results were subsequently compared for each type of proxy to reconstruct hydroclimate patterns and, in a second step compared to simulation output derived from the Climate Com-munity System Model version 3 (CCSM3). Reconstructed P-E and precipitation were assessed by the following procedures. Pollen re-cords were biomized following the approach of Prentice et al. (1996), which is based on the relationship between plant functional types (PFTs) and biomes. These relationships have been established for China (Yu et al., 2000a), Japan (Takahara et al., 2000), Southeast Asia, (Pickett et al., 2004) and the IOM region (Kramer et al., 2010). Although affinity scores were not calcu-lated here, qualitative P-E was directly approximated from the biomes based on the relation-ship between biomes and mean annual precipitation and temperature, as suggested by Mader (2012). Inferred P-E was grouped into three categories, representing low (1), medium (2) and high (3) (see Table 3). For the temporal analysis, the expansion/shrinkage of grassland was assessed using non-arboreal pollen taxa (e.g. Artemisia, Compositae, Rosaceae, Chenopodi-

19

aceae, Poaceae) and higher/lower run-off was reconstructed using changes in other major eco-logical groups (e.g., Pteridophyta spores). Changes in grassland and run off were assumed to reflect local and regional hydroclimate changes.

Fig. 6: Location of marine and terrestrial records included in the compilation of Paper I. See Appendix I for details on each record.

Fig. 7: Location of published pollen-stratigraphic records included in Paper II. See details for each record in Appendix II.

For the speleothem records, we assumed that the speleothems grew in a closed system (∆T ~0°C) and that variations in δ18O reflect the precipitation amount over the site (Wang et al., 2001; Yuan et al., 2004). The mean LGM precipitation was reconstructed based on average δ18O values between 23 and 19 ka BP for each site following the categorization in Table 3. The

20

speleothem δ18O values were then smoothed using a five-point running mean to provide in-sight into Asian summer monsoon variability between 25 and 17 ka BP. δ18O records of Globigerinoides ruber (G. ruber) were selected from marine stratigraphies. For this, we assumed that (i) LGM sea level was stable (Clark and Mix, 2002; Clark et al., 2009), (ii) major change in salinity did not occur during the LGM (Schulz et al., 1998; Clark and Mix, 2002), and (iii) the δ18O values of LGM global ocean effect were constant ~1.1‰ higher than present (Adkin et al., 2002; Ravelo and Hillaire-Marcel, 2007). SST differences were calculated from the MARGO (2009) and the World Ocean Atlas (WOA) 1998 (Conkright et al., 1998) data sets, and converted to δ18O values using the relationship of a 0.25‰ increase in δ18O with a SST decrease of 1°C (Erez and Luz, 1983). P-E was calculated by subtracting the δ18O values of the LGM global ocean effect and the SST difference from the δ18O values of G. ruber. The ob-tained P-E values were then smoothed using a five-point running mean between 25 and 17 ka BP. Table 3: Climate variables used to assess P-E and precipitation in Paper I.

Effective moisture Palaeo-proxy assessment Biome P-E

Qualitative precipitation Speleothem δ18O values

3

Warm-temperate rain forest, tropical deciduous broadleaf forest and wood land, wet sclerophyll forest, and broadleaved evergreen/warm forest

<-1.4 <-7.70

2 Mix of steppe and warm mixed forest

1 Cool mixed forest, mix of steppe and cool mixed forest and steppe

>0.3 >-0.58

Bulk sediment δ13C, hardwood and faunal remains were used as supporting information fol-lowing the original interpretation of the respective authors. Climate model output data

The Community Climate System Model version 3 (CCSM3) is a fully coupled global atmos-phere-ocean-sea-ice-land surface climate model (Collins et al., 2006). CCSM3 output compares well with palaeo-data in other studies about extreme climate conditions (e.g. Kjellström et al., 2009; Brandefelt et al., 2011). The CCSM3 simulation used here is described in detail in Otto-Bliesner et al. (2006) and Brandefelt and Otto-Bliesner (2009). CCSM3 simulated precipitation and P-E were compared with the qualitative data set of the compiled palaeo-proxy data. The mean position of the ITCZ was further calculated based on the simulated precipitation accord-ing to the method outlined by Braconnot et al. (2007) as:

∑

∑

=

==

N

prlaty

N

prlaty

ypr

ylatypr

lonITCZloco

o

30

max)(

30

max)(

)(

)()(

)(_

where )(_ lonITCZloc = mean position of the ITCZ in each longitude

21

max)( prlat = latitude of maximum precipitation in each longitude

)( ypr = precipitation over latitude y

2.1.2 Reconstruction of semi-quantitative precipitation and temperature in Paper II

The pollen diagrams were digitized and the biomization followed Prentice et al. (1996). The biomization is based on the relationship between PFTs and the biomes that had been estab-lished by Picket et al. (2004). The affinity scores were calculated as:

[ ]∑ −=i jkjjiki pA θδ

,,,

where Ai,k

represents the affinity score of biome i in sample k .

ji ,δ is entry of taxon j in biome i based on taxon×biome matrix (if taxon j is included in

biome i ,ji ,

δ = 1, otherwise ji ,

δ = 0).

kjp,

is the percentage of pollen taxon j in sediment sample k .

jθ denotes the threshold of pollen taxon j which here is 0.5% in order to eliminate the

influence of long-transported pollen (e.g. Prentice et al., 1996; Herzschuh et al., 2006).

Affinity score percentages of each biome were subsequently calculated and analysed together based on the assumption that all pollen taxa contained in the sediment samples had been di-rectly or indirectly influenced by hydroclimatic variations. Semi-quantitative precipitation and temperature were separately assessed using the Southeast Asian and Pacific (SEAPAC) model of biome distribution in climate space (Pickett et al., 2004) (Table 4). Precipitation ― The precipitation range of each time interval (

tP ) was assumed to be a func-

tion of precipitation (iP ) and affinity score percentage (

ix ) of each biome (see Table 4)

(Picket et al., 2004):

3

3

2

2

1

1

100100100P

xP

xP

xPt

+

+

= ( )321 PPP <<

where

1P is the precipitation of temperate or tropical grassland and xerophytic shrubland

(STEP), and of xerophytic woods/scrub (XERO) biomes. We here assume that 1P is less than

500 mm a-1 (Pickett et al., 2004).

2P represents the precipitation of tropical deciduous broadleaf forest and woodland

(TDFO), and of temperate sclerophyll woodland and shrubland (DSFW) biomes and is esti-mated to between ~500 and ~1100 mm a-1 (Pickett et al., 2004).

3P denotes the precipitation of tropical semi-evergreen broadleaf forest (TSFO), wet

sclerophyll forest (WSFW), tropical evergreen broadleaf forest (TRFO), warm-temperate rain-forest (WTRF), cool-temperate rainforest (CTRF) and cool evergreen needle-leaf forest (COCO) biomes.

3P is here assumed to be higher than ~1100 mm a-1 (Pickett et al., 2004)

1x , 2x and x

3are the sums of the percentages of the affinity scores of STEP and XERO,

of TDFO and DSFW, and of TSFO, WSFW, TRFO, WTRF, CTRF and COCO, respectively;

XEROSTEP xxx +=1

DSFWTDFO xxx +=2

COCOCTRFWTRFTRFOWSFWTSFO xxxxxxx +++++=3

22

We further divided

3P into 13−P and

23−P which represent the precipitation of TSFO and

WSFW, and TRFO, WTRF, CTRF and COCO biomes, respectively. The moisture index (α) be-tween

13−P and 23−P is 0.8 and

13−P is higher than 23−P (Picket et al., 2004).

13−x and 23−x are the

affinity score percentages of 13−P and

23−P , respectively. The semi-quantitative precipitation

was approximated from 13−x and

23−x ratio, which deviate from the moisture index of 0.8.

Table 4: Conceptual model of biome distribution in climate space (modified from Picket et al., 2004).

Increasing moisture index (α) P~500 mm a-1 P~1100 mm a-1 α:0.8

Tropical deciduous broadleaf forest and

woodland (TDFO)

Tropical semi-evergreen broadleaf

forest (TSFO)

Tropical evergreen broadleaf forest

(TRFO)

18°C

Warm-temperate rainforest (WTRF)

Frost limit

Cool-temperate rain-forest (CTRF)

5°C

Incr

easi

ng

mea

n t

emp

erat

ure

of t

he c

old

est

mon

th

Tem

per

ate

or t

rop

ical

gra

ssla

nd

an

d x

erop

hyti

c sh

rubl

and

(S

TE

P)

Xer

ophy

tic

woo

ds/

scru

b (X

ER

O)

Tem

per

ate

scle

rop

hyll

woo

dla

nd

an

d s

hru

blan

d

(DSF

W)

Wet

scl

erop

hyll

fore

st (

WSF

W)

Cool evergreen nee-dle-leaf forest

(COCO)

Tundra (TUND)

Temperature ― Semi-quantitative temperatures at each time interval (

tT ) were estimated fol-

lowing the same assumption and procedure as for precipitation. STEP and XERO biomes were here excluded because of their distribution in all climate spaces (Picket et al., 2004) (Table 4).

2

2

1

1

100100T

yT

yTt

+

= ( )21 TT <

where

1T and 2T represent the temperatures of the DSFW, WSFW, WTRF, CTRF and COCO

biomes, and of the TDFO, TSFO and TRFO biomes, respectively.

1y and 2y are the percentages of the affinity score of

1T and 2T .

23

COCOCTRFWTRFWSFWDSFW yyyyyy ++++=1

TRFOTSFOTDFO yyyy ++=2

1T is less than2T which is higher than 18°C. Semi-quantitative temperature was assessed by

2y and1y which deviated from 18°C.

2.2 Lake sediment analysis Sediment cores from Lakes Kumphawapi and Ph Kho in northeast Thailand were collected during fieldwork in January 2010 and 2011 using a modified Russian corer (7.5 cm, diameter) (sites #1 and 2 in Fig. 5 and Fig. 8). To achieve a continuous sequence, we collected each one-meter sequence with an overlap of 50 cm.

The sediment cores were transported to the Department of Geological Sciences, Stockholm University and stored in a cold room prior to sub-sampling and laboratory analyses.

The sediment cores were carefully cleaned prior to sub-sampling and the stratigraphy of each sediment core was described in detail. Specific marker horizons were used to correlate between overlapping core segments to create a composite sequence for each site.

Each 1-m long sediment core was subsequently scanned every 5 mm using an Itrax XRF core scanner with a Mo tube (30 kV and 30 mA for measuring 60 sec/point). All elements obtained from XRF scanning were smoothed using a 3-point running mean to create 1-cm resolution profiles. Some selected elements were normalized by incoherent (Compton) and coherent (Rayleigh) scattering and then used in a statistical approach to assess their relationship and to reconstruct environmental changes. Consecutive 1-cm samples were dried overnight at 105°C to measure loss-on-ignition (LOI). The dried samples were combusted for 6 hr at 550°C and for 4 hr at 950°C, respectively. The weight loss of the samples combusted at 550 °C and 950°C, respectively was calculated in re-spect to the dried sample weight and is reported as % LOI.

Selected sediment samples were freeze-dried and homogenized prior to analyses of total organic carbon (TOC), nitrogen (TN) and sulphur (TS), and their isotopic composition (δ13C, δ15N and δ34S). These proxies were analysed using a Carlo Erba NC2500 elemental analyzer connected to a Finnigan MAT Delta+ mass spectrometer. δ13Corg, δ15N and δ34S were measured relative to the Vienna PeeDee Belemnite (VPDB), AIR and Canyon Diablo Troilite (CDT) stan-dard, respectively. Analytical errors for δ13Corg and δ15N were ±0.15‰ and for δ34S measure-ments the uncertainty was ±0.2‰. The carbon:nitrogen ratio (C/N) was multiplied by 1.167 to yield atomic ratios. For biogenic silica (BSi), sediment samples were pre-cleaned with H2O2 and HCl to remove organic matter and carbonate. The cleaned samples were then extracted by 40 mL of 1% Na2CO3 solution for 5 hr and neutralized with 0.21N HCl. The dissolved silica concentration was analysed after the extraction by using ICP-OES (Varian Vista Ax) and corrected by a simul-taneous dissolution of minerogenic silica to determine weight percentages (wt %) of BSi. Sediment samples were sieved under running tap water by using sieves with a mesh size of 0.5 mm. The sieved plant remains were stored in deionized water. The plant remains were carefully identified under a stereomicroscope, cleaned multiple times with deionized water, and dried overnight at 105°C in glass vials. Selected samples were sent to the 14CHRONO Cen-tre at Queen’s University, Belfast for radiocarbon analysis. The radiocarbon age and one stan-dard deviation were calculated following the conventions of Stuiver and Polach (1977) using the Libby half-life of 5568 years and a fractionation correction based on δ13C measured on the AMS. The age-models were constructed using Bacon, a Bayesian statistics-based routine that models accumulation rates by dividing a sequence into many thin segments and estimating the (linear) accumulation rate for each segment based on the (calibrated) 14C dates, together with

24

assumptions about accumulation rate and its variability between neighbouring segments (Blaauw and Christen, 2011).

Fig. 8: (A) Location of Lakes Kumphawapi and Pa Kho, and (B) topography around the lakes and coring points. Sediment cores from Lake Kumphawapi (core KMP-CP4) and Lake Pa Kho (core PK-CP3) were used for Papers III and IV.

25

3. Summary of the papers Below I outline my contributions to the analysis, data interpretation and writing of the four

papers included in this thesis (Table 5).

Table 5: Author contribution to the analysis, data interpretation and writing of the four papers included in this thesis.

Paper I Paper II Paper III Paper IV Palaeo-data compi-lation

A. Chabangborn A. Chabangborn

Climate model (CCSM3) simulation

J. Brandefelt

Coring and field-work

A. Chabangborn B. Wohlfarth L. Löwemark S. Fritz S. Inthongkeaw W. Klubseang

A. Chabangborn B. Wohlfarth L. Löwemark S. Fritz S. Inthongkeaw W. Klubseang

Lithostratigraphy A. Chabangborn B. Wohlfarth

A. Chabangborn B. Wohlfarth S. Chawchai

Loss-on-ignition A. Chabangborn A. Chabangborn ITRAX core scan-ning

A. Chabangborn L. Löwemark

A. Chabangborn L. Löwemark

Carbon, nitrogen and sulphur elemen-tal and isotope analysis

A. Chabangborn A. Chabangborn S. Chawchai

Biogenic silica S. Chawchai S. Chawchai Radiocarbon sam-ples and dating

P.J. Reimer B. Wohlfarth

P.J. Reimer B. Wohlfarth

Age-depth model construction

M. Blaauw S. Chawchai M. Blaauw

Figures and tables A. Chabangborn A. Chabangborn S. Chawchai S. Chawchai A. Chabangborn (Fig. 5)

Data interpretation A. Chabangborn J. Brandefelt

A. Chabangborn S. Chawchai B. Wohlfarth A. Chabangborn

S. Chawchai A. Chabangborn B. Wohlfarth

Manuscript writing A. Chabangborn A. Chabangborn B. Wohlfarth

S. Chawchai B. Wohlfarth

S. Chawchai

26

4. Results The results from the three published papers and one submitted manuscript included in this

thesis are summarized below.

4.1 Paper I: Asian monsoon climate during the Last Glacial Maximum: pa-

laeo-datamodel comparisons The Asian monsoon region was separated into three sub-regions: the IOM, EAM, and Sun-

daland and Sahulland (SSS) (Fig. 9). The IOM and EAM sub-regions were divided along longi-tude 105°E following Wang et al. (2003) and Wang et al. (2005). The LGM sea level low stand exposed the vast area of the Sunda and Sahul continental shelves, which are included as a third domain (Bird et al., 2005; De Deckker et al., 2002; Hanebuth et al., 2011) (Figs. 9A, 9B).

Palaeo-data and model comparison

Indian Ocean monsoon sub-region – Enrichment in δ18O values of speleothems from the western Arabian Sea (site #33) (Shakun et al., 2007) implies low precipitation, which agrees well with the CCSM3 simulation (<1000 mm a-1) during the LGM (Table 3) (Fig. 9A). Reconstructed P-E based on planktonic foraminifera obtained from the northern Arabian Sea is low (site #34) (van Rad et al., 1999), while high P-E was reconstructed by the same proxy off south India (site #29) (Tiwari et al., 2006). These wet and dry conditions are consistent with the model output (Fig. 9B).

Simulations of moderately wet conditions (0-800 mm a-1) compare moderately well with re-constructions of high P-E based on δ18O values of planktonic foraminifera in marine records from the Bay of Bengal (site #32) (Rashid et al., 2007) (Fig. 9B).

Data-model discrepancies are visible in Sri Lanka (site #30) (Premathilake and Risberg, 2003; Premathilake, 2006), southern India (site #31) (Rajagopalan et al., 1997), and in the northern part of the IOM sub-region (site #35) (Sharma and Chatterjee, 2007); (site #36) (Kot-lia et al., 2010); (site #37) (Fujii and Sakai 2002; Hayashi et al., 2009); (site #38) (Cook et al., 2011); (site #39) (Tang et al., 1998; Tang et al., 2004); (site #40) (Wattanapituksakul, 2006) (Figs. 9A, 9B). CCSM3 simulation output shows moderate P-E, while pollen records from Sri Lanka (site #30) (Premathilake and Risberg, 2003; Premathilake, 2006) and the northern IOM sub-region (site #36) (Kotlia et al., 2010); (site #37) (Fujii and Sakai 2002; Hayashi et al., 2009); (site #38) (Cook et al., 2011); (site #39) (Tang et al., 1998; Tang et al., 2004), a distinct barren pollen zone from northern India (site #35) (Sharma and Chatterjee, 2007) and a domi-nance of dry-tolerant C4 plants in sediments from the Nilgiri Hills (site #53) (Rajagopalan et

al., 1997) indicate arid conditions during the LGM (Figs. 9A, 9B). Reconstructions of wet condi-tions based on faunal remains from west Thailand (site #40) (Wattanapituksakul, 2006) also diverge from the CCSM3 simulations of moderately wet conditions.

Sundaland and Sahulland sub-region - Reconstructed wet conditions correspond well with simulated high precipitation (>2000 mm a-1) and P-E (>800 mm a-1) over the SSS sub-region during the LGM (Figs. 9A, 9B). Pollen assemblages were assigned to warm-temperate rain for-est and tropical deciduous broadleaf forest and woodland biomes in western (site #7) (Newsome and Flenley, 1988); (site #9) (Maloney and McCormac, 1996); (site #8) (Maloney, 1980), southern (site #2) (Stuijts et al., 1988); (site #3) (van der Kaars and Dam, 1997); (site #4) (van der Kaars et al., 2010), and central (site #6) (Anshari et al., 2001) Sundaland; from the northern Molucca Sea region (site #10) (Barmawidjaja et al., 1993), the Banda Sea (site #5) (van der Kaars et al., 2000) and the eastern part of the SSS sub-region (site #1) (Hope, 2009), which implies high P-E (Table 3) (Fig. 9B). The plant remains of Dipterocarpaceae, which can be assigned to warm-temperate rain forests (Pickett at al., 2004), also indicate high P-E over the Malay Peninsula (site #11) (Wüst and Bustin, 2004) during the LGM. In addition, wet con-

27

ditions are consistent with low δ18O values of speleothems from northern Borneo (site #12) (Partin et al., 2007) (Table 3).

Fig. 9: Comparison between qualitative reconstruction and quantitative simulation of (A) pre-cipitation and (B) P-E during the LGM. The CCSM3 precipitation was classified as low (<1000 mm a-1), medium (1000-2000 mm a-1) and high (>2000 mm a-1). Simulated P-E was also cate-gorized as dry (<0 mm a-1), moderately wet (0–800 mm a-1), and wet (>800 mm a-1) conditions (modified from Chabangborn et al., 2013).

East Asian monsoon sub-region - Pollen assemblages from Southern China (site #16) (Zheng and Lei, 1999) (site #17) (Wang et al., 2010c) were assigned to a mix of warm mixed forest and steppe biomes and indicate moderate P-E during the LGM (Table 3). While in Japan (site #26) (Hayashi et al., 2010); (site #27) (Takahara and Takeoka, 1992); (site #28) (Yasuda, 1982) the assemblages were assigned to and a mix of steppe and cool mixed forest biomes suggesting low P-E, (Table 3). These hydroclimate conditions compare well with CCSM3 model output of

28

moderate P-E (0-800 mm a-1) over the southern part of China and low P-E (>800 mm a-1) over Japan (Fig. 9B).

Reconstructions of low P-E based on pollen stratigraphies from Korea (site #25) (Chung et

al., 2006) and from the Chinese loess plateau (site #23) (Sun et al., 1997); (site #24) (Wang and Sun, 1994) generally disagree with simulations of moderate P-E (Fig. 9B). The recon-structed palaeo-vegetation from Taiwan (site #18) (Liew et al., 2006) indicates high P-E which markedly diverges from simulations of low P-E. Data-model discrepancies are also found in the reconstructed precipitation based on δ18O values of speleothems from Jintawan Cave (site #20) (Cosford et al., 2010) and Hulu Cave (site #21) (Wang et al., 2001) (Fig. 9A). The low δ18O values of the speleothems imply wet conditions, while CCSM3 simulates moderately wet condi-tions during the LGM (Fig. 9A).

Fig. 10: (A) Qualitative precipitation change inferred from biomes and δ18O of speleothems in the Asian monsoon region between 25 and 17 ka BP. Thick orange and blue lines represent the calibrated age points and the thin orange and blue lines are interpolations between these. The blue and orange colours represent relatively wet and dry climatic condition, respectively. (B) The map shows the location of the sites included in the temporal variability analysis (modi-fied from Chabangborn et al., 2013). See Appendix I for details on the sites.

29

Temporal variability

Temporal changes in hydroclimate patterns were assessed using palaeo-records with at least two dating estimates between 25 and 17 ka BP. In the IOM sub-system, most palaeo-records indicate that hydroclimate conditions shifted from dry to wet between 20 and 19 ka BP (site #30) (Prematilake and Risberg, 2003; Prematilake, 2006); (site #33) (Shakun et al., 2007); (site #35) (Kotlia et al., 2010); (site #36) (Fujii and Sakai, 2002; Hayashi et al., 2009); (site #37) (Cook et al., 2011) (Fig. 10).

A pollen stratigraphy from western Sumatra (site #7) (Newsome and Flenley, 1988) and δ18O values of speleothems from northern Borneo (site #12) (Partin et al., 2007) indicate wet-ter conditions ~20 ka BP and drier conditions thereafter (Fig. 10). In contrast, a decrease in P-E was reconstructed by pollen records from the Sumatra plateau for approximately the same time interval (site #9) (Maloney, 1980); (site #7) (Maloney and McCormac, 1996) (Fig. 10). Pollen records from South China (site #17) (Wang et al., 2010c) and Taiwan (site #18) (Liew et al., 2006) depict drier conditions around ~20-19 ka BP (Fig. 10). On the other hand, the decrease in δ18O values of speleothems from sites around 30°N (site #34) (Cosford et al., 2010); (site #35) (Wang et al., 2001); (site #36) (Zhou et al., 2008) imply an increase in pre-cipitation between ~20 and ~19 ka BP. These wetter conditions seem consistent with a recon-structed increase in P-E based on pollen records from Japan (site #42) (Yasuda, 1982). Other pollen records from Japan (site #26) (Hayashi et al., 2010) however do not show any hydro-climate change throughout the LGM. 4.2 Paper II: Climate over mainland Southeast Asia 10.5-5 ka BP

The number of data points available for this reconstruction varied greatly (Fig. 11A). No data were available for the interval between 12.5-11.5 ka BP and the few data points between 15-12.5 ka BP were excluded from the published paper. However, there are 12-17 data points for 8.5-7 ka BP, and 3-5 data points for 10.5-8 ka BP and for 7-5 ka BP.

Pollen assemblages were mainly assigned to warm-temperate rainforest (WTRF), tropical deciduous broadleaf forest and woodland (TDFO), tropical semi-evergreen broadleaf forest (TSFO), and tropical evergreen broadleaf forest (TRFO) biomes (Fig. 11B). These biomes indi-cate that precipitation and temperature generally exceeded 1100 mm a-1 and 18°C over MSEA between 15 and 5 ka BP (Picket et al., 2004). Semi-quantitative precipitation and temperature variability

The dominance of the WTRF biome (WTRFx >50%) implies cool and wet conditions from 15

to 13.5 ka BP (Figs. 11B, 11C). WTRFx abruptly decreases at ~13 ka BP when

TDFOx and

TSFOx rapidly increase. These changes in pollen assemblages suggest a climate shift from

cool/wet to warm/dry climate conditions over MSEA ~13 ka BP (Fig. 11C).

TDFOx is almost constant ~40%, while WTRFx varies by 23-44% from 11.5 to 5 ka BP (Fig.

11B). TSFOx

increases from 13 to 22% ~10 ka BP and varies about 20% thereafter. TRFOx in-

creases from 5 to 10% between 11.5 and 10 ka BP and gradually decreases from 10 to 5 ka BP. The biome assignments thus indicate that precipitation decreased between 11.5 and 10.5 ka BP. The amount of precipitation seems to have been similar at 10.5 and at 7 ka BP and gener-ally declined after 7 ka BP (Fig. 11C). However, this precipitation pattern was punctuated by dry intervals between 10-9 ka BP and 6.5 ka BP. Despite insignificant changes in temperature, warmer intervals coincide with drought periods (Fig. 11C).

30

Fig. 11: (A) Number of pollen samples included in the reconstruction; (B) variability of the ma-jor biomes; and (C) reconstructed semi-quantitative precipitation and temperature between 15 and 5 ka BP. Coloured lines represent the calibrated age points of each sample. Filled squares represent published data sets and open squares unpublished data set. See Appendix 2 for details on the sites (modified from Chabangborn and Wohlfarth, 2014). Spatial analysis

Precipitation and temperature over MSEA during the early (10.5-8.0 ka BP) and mid Holo-cene (8.0-5.0 ka BP) were reconstructed based on the semi-quantitative data set (Fig. 11). The inferred biomes suggest wet to moderately wet conditions over MSEA during the early Holo-cene (Fig. 12A). Reconstructed dry conditions in site #3 (Kealhofer and Penny, 1998) and #5

31

(Li et al., 2012) however are possibly caused by the influence of local conditions (e.g. catch-ments, soil type and/or human disturbances), or due to the low diversity pollen spectra. High to moderate precipitation during the mid Holocene was reconstructed for sites located south of 12°N (Fig. 12B). The reconstruction of wet conditions over site #4 (Penny, 2001) may come from an underestimation of the non-arboreal pollen assemblages as discussed by Burke (2014). Pollen records suggest warm to moderately warm conditions over MSEA between the early and mid Holocene (Figs. 12C, 12D). Reconstructed cool conditions at site #13 (Rugmai, 2006) may also be due to the low diversity pollen spectra.

Fig. 12: (A, B) Precipitation and (C, D) temperature over MSEA during the early (10.5-8 ka BP) and mid Holocene (8-5 ka BP). The grey land areas show the palaeo-shoreline after the sea level rise at 10.5 ka BP and for the sea level high stand at ~6 ka BP (Hanebuth et al., 2011; Chabangborn and Wohlfarth, 2014).

32

4.3 Paper III: Lake Kumphawapi – an archive of Holocene palaeoenviron-

mental and palaeoclimatic changes in northeast Thailand Several lake sediment cores and a variety of proxies were analysed from Lake Kumphawapi,

Thailand (Figs. 5, 8) to reconstruct hydroclimatic changes during the Holocene. Sediment se-quence KMP-CP4, which is the focus of this paper, was collected from the southern part of the lake. The age-depth model suggests continuous sediment and peat deposition until 6 ka BP, a hiatus of several thousand years, and sediment deposition since 1.6 ka BP.

The lowermost sediments (units 1a, b) between 7.00 and 3.80 m date to 9.8-7.6 ka BP and are made up of silty gyttja clay and gyttja clay with low %LOI and %TOC (units 1a, b). This suggests low lake organic productivity (Fig. 13). The increase in Zr/Rb ~6.00-5.50 m, which indicates coarser grain sizes (Dypvik and Harris, 2001), coincides with an increase in the C/N ratio (Figs. 13, 14C). This could point to higher run off to the lake between 9.80 and 9.60 ka BP and in-wash of terrestrial organic matter. Si, Ti, Rb, and Ca values are well correlated with each other between 5.40-3.80 m, which would suggest transport of clay- and silt-sized particles to the lake and relatively stable lake conditions.

The organic matter content increases slightly in the gyttja clay of unit 2a (Fig. 13). The geo-chemical proxies suggest a stable lake with increasing aquatic productivity and less run off between 7.6 and 7 ka BP.

Fig. 13: Lithostratigraphy, %LOI, %TOC, C/N ratios, %BSi and δ13C values of core KMP-CP4 (Chawchai et al., 2013).

33

The transition between the gyttja clay in unit 2a and the peaty gyttja and peat in unit 2b to-gether with changes in the geochemical proxies suggest that the lake developed into a wetland and eventually a peatland. Peaks of Si, Zr, Ti and Zr/Rb in the peat of unit 2b compare well with low %LOI and %TOC between 2.91 and 2.93 m (Figs. 13, 14B) and might be attributed to wind transport under dry conditions. The lithological and geochemical proxies thus indicate a grad-ual decrease in P-E between 7 to 6.5 ka BP. The hiatus, which is present between 6 and 1.6 ka BP may also relate to dry conditions and to no or low sedimentation/peat growth.

Fig.14 : XRF core scanning profiles of major elements (Si, Zr, K, Rb, and Ca), and Zr/Rb ratio of (A) units 3a-3b, (B) unit 2b and (C) unit 1a-2a (modified from Chawchai et al., 2013).

The deposition of clayey gyttja and algae gyttja in units 3a and 3b suggests the re-establishment of a shallow lake and possibly an increase in P-E after 1.6 ka BP (Fig. 13). The increase in %LOI and %TOC and the decrease in the C/N ratio suggest an increase in the lake

34

organic productivity and a predominance of aquatic sourced organic matter ~0.6 ka BP (Fig. 13). All major elements (Si, Zr, K, Ti, Rb, and Ca) decrease at 2.50 m, which would indicate lower run off into the lake (Fig. 14A). However, the sediments and the geochemical proxies do not show signals of higher run off. This discrepancy may be explained by the influence of hu-man activities (higher nutrient transport due to agriculture) after 0.6 ka BP, since archaeologi-cal evidence suggests a settlement at Ban Don Kaew ~1.1 ka BP (Fig. 8).

4.4 Paper IV: Hydroclimatic shifts in northeast Thailand during the last two

millennia – the record of Lake Pa Kho The sediment sequence used in Paper IV is derived from Lake Pa Kho, which is located ~15 km to the southwest of Lake Kumphawapi (Fig. 8). Paper IV focused on lake status changes and hydroclimatic conditions based on multi-proxy analyses of the sediment sequence between 2 to 3.5 m, corresponding to BC 170 - AD 2001.

Lithological units 5 (3.50-3.36 m) and 4 (3.36-3.34 m) consist of fine detritus gyttja and peaty gyttja, respectively. These sedimentary units are characterised by a slight change in C/N ratios (27-24) and by a gradual decline in δ13C values (-24 to -21‰), which imply a mix of ter-restrial and aquatic sources of the organic matter (Fig. 15). This is consistent with the presence of both phytolith and diatoms, and of terrestrial, telmatic and aquatic species in the plant mac-rofossil record. The elevated values of Si, Ti and Ca suggest that run off was important, which is also supported by the terrestrial plant input. These evidences indicate that Lake Pa Kho was a shallow water lake, which received run off from the surrounding catchment between BC 170 and AD 370. The lithological transition from peaty gyttja (unit 4) to compact peat (unit 3: 3.34-2.73 m) at ~AD 370 is marked by a decrease in δ13C values, and the presence of phytoliths and of terres-trial plant remains (Fig. 15). δ13Corg values increase again between 2.98 and 2.93 m (AD 410-450) and fluctuate between -30 and - 27‰ until 2.68 m (AD 800). Despite relatively constant %TOC, the gradual decline in C/N ratios from 2.93 to 2.68 m can be potentially explained by the consumption of organic carbon during peat decomposition (Chimner and Ewel, 2005; Ise et

al., 2008). High amounts of charcoal can be observed between 2.98 and 2.73 m (AD 410-650). Plant remains change from terrestrial species at the bottom to terrestrial/telmatic species at the top of unit 3. The multi-proxies thus indicate low P-E from AD 370 to 800, which was how-ever punctuated by a short interval of higher P-E between AD 410 and 450.

The re-establishment of a wetland and higher P-E may be inferred from the abrupt increase in δ13Corg values and the shift from terrestrial/telmatic plant species to telmatic/aquatic plant species between AD 800 and 970 (Fig. 15). Although the age model suggests a 330-year long hiatus at 2.63 m, the stratigraphy does not give any indications for a break in peat growth. Therefore we present two alternative hypotheses: one including a hiatus, and one inferring slow peat growth.

The δ13Corg values of -28‰ indicate deposition of predominantly terrestrial plant remains between 2.63-2.62 m, which would suggest lower P-E between AD 1300-1340. This interval is followed by a gradual increase in δ13Corg values, by the presence of diatoms and a shift from telmatic to aquatic plant species, which all signify wetter conditions after AD 1450. The step-wise increase in Si, K, and Ti at 2.17 m depth (AD 1850) and at 2.05 m depth (AD 1970) could be related to higher run-off due to land-use changes.

35

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5. Discussion 5.1 The last glacial maximum (23-19 ka BP): Paper I

Reconstructed P-E and precipitation agree well with CCSM3 output, which implies wet con-ditions in the SSS domain as well as off southern India, moderately wet conditions in south China, and dry conditions in the west Arabian Sea (Figs. 9A, 9B). Reconstructions of wet condi-tions over the Bay of Bengal compares fairly well with CCSM3 simulations of moderately wet conditions. Data-model discrepancies are observed in Sri Lanka, southern India and north of latitude 30°N, where palaeo-proxies suggest dry conditions, but CCSM3 suggests wet condi-tions (Fig. 9B). Differences between reconstructions and simulations are also found in north-west Thailand and central China, where palaeo-proxies provide wetter conditions than those simulated by CCSM3 (Fig. 9A).

The data–model discrepancies can be caused by the palaeo-proxy assessment and the CCSM3 simulation. The qualitative data assessed by palaeo-proxies are difficult to categorize with a proper method. The reconstructed conditions consequently may not have been as dry or wet as here assumed. Moreover, the proxies are possibly showing more response to other en-vironmental and climatic parameters other than precipitation or P-E. The δ18O values of spe-leothem records from China, for example, are generally interpreted as precipitation amount (e.g. Wang et al., 2001; Cosford et al., 2010), while recent studies challenge this and argue that they possibly reflect changes in atmospheric circulation (Clemens et al., 2010; Pausata et al., 2011; Maher and Thomson, 2012). On the other hand, the low-resolution of CCSM3 simulation and the selected LGM boundary conditions may not well capture ocean-atmosphere interac-tions and other control parameters, e.g. LGM dust (Otto-Bliesner et al., 2006; Marchant et al., 2007).

A weakening of the summer monsoon over the IOM and the northern EAM sub-regions was inferred from reconstructed dry conditions, which could be explained by a decrease in land-sea thermal contrast caused by a slight change in palaeo-shorelines and a decrease in sea sur-face temperatures (Yokoyama et al., 2000; MARGO, 2009; Hanebuth et al., 2009, 2011). Aridity may have been caused by reduced transport of humid air masses to the western part of the IOM sub-region, and a strengthening of the Siberian high-pressure system and westerly winds in the northern part of EAM sub-region (An et al., 2012). On the other hand, reconstructed wet conditions in the southern part of the EAM domain and in the SSS sub-area, where a vast area was exposed by the LGM sea level low stand, indicate a strengthening of the summer monsoon. The stronger summer monsoon intensity in the SSS sub-region and in south China can be ex-plained by an increase in land-sea thermal contrast cause by decreased SSTs and increased land areas (De Deckker et al., 2002; Bird et al., 2005). All these together indicate that the LGM mean position of the ITCZ was located off southern India and in the southern part of China (Fig. 16). Reconstructions of the mean position of the ITCZ correspond well with CCSM3 output in the west Arabian Sea, and compare moderately well over MSEA. Palaeo-data – model discrep-ancies are observed in southern India and in the Bay of Bengal, where the CCSM3 simulated ITCZ shifts northward, while the palaeo-records suggest dry conditions.

A distinct hydroclimate reversal, potentially related to the northward shift of the mean posi-tion of the ITCZ, is reconstructed from wetter conditions in the northern part of the EAM sub-region and in the IOM domain (Figs. 10, 16, 17). A strengthening summer monsoon and a northward shift in the mean position of the ITCZ was possibly caused by the rapid sea level rise, which led to the reorganization of the atmospheric circulation at ~19 ka BP (Hanebuth et

al., 2011) (Fig. 4).

37

Fig 16: Reconstruction of the mean position of the ITCZ (dashed red line) in the Asian monsoon region during the last glacial maximum (LGM) (23-19 ka BP) and 19.5 ka BP (modified from Chabangborn et al., 2013). The green lines represent the seasonal shift of the present-day ITCZ according to Wang (2009). The grey land areas mark the palaeo-shorelines during the LGM sea level low stand (-120 m) and after the rapid sea level rise ~19.5 ka BP following Hanebuth et

al. (2011). 5.2 The early to mid Holocene (10.5-5 ka BP): Paper II and III

The semi-quantitative data set indicates that precipitation and temperature generally ex-ceed 1100 mm a-1 and 18°C, respectively, over MSEA between 15 and 5 ka BP. In addition, the data set suggests that cool/wet and warm/dry conditions alternated. This assumption diverges from the general assumption of cool/dry and warm/wet conditions (e.g. Mingram et al., 2004; Herzschuh et al., 2009; Zhang and Mischke, 2009). The association of cool/wet and warm/dry conditions can be explained by the importance of cloud cover that plays an important role in Earth’s energy budget in this area (Trenberth and Shea, 2005). A decrease in the amount of cloud cover due to lower humidity can enhance solar irradiance and thus causes warmer sur-face temperature, and vice versa. The reconstruction of high precipitation over MSEA (unpublished data sets) compares well with speleothem records from the IOM and EAM sub-systems (Sinha et al., 2005; Dykoski et al., 2006; Partin et al., 2007; Shakun et al., 2007; Yang et al., 2010), which indicate a strengthening of the summer monsoon over the entire Asian monsoon region between 15 and 13 ka BP (Figs,

38

11, 17). The interval of intense summer monsoon coincides with warmer conditions in the Northern Hemisphere, and concomitant cooler conditions in the Southern Hemisphere, as seen from Greenland and Antarctic ice core records (NGRIP Members, 2004; EPICA Community Members, 2006). The strengthening of the summer monsoon can be explained by a northward shift of the mean position of the ITCZ, which was caused by the Northern and Southern Hemi-sphere temperature gradient (Linzen and Nigam, 1987; Lindzen and Hou, 1988; Chiang and Bitz, 2005; Kang et al., 2009; Cvijanovic, 2013).

Fig. 17: Summary of hydroclimate variability over mainland Southeast Asia from 25 ka BP to present day, based on the results of Papers I, II, III and IV. Drier conditions associated with warmer temperature over MSEA occurred between 13 and 12.5 ka BP. This is slightly later than the decrease in precipitation reconstructed for Oman (Shakun et al., 2007) and Borneo (Partin et al., 2007), but earlier than that observed for India (Sinha et al., 2005) and China (Dykoski et al., 2006; Yang et al., 2010) (Figs, 11, 17). The inter-val of a weaker Asian summer monsoon corresponds with the Younger Dryas cooling in north-ern hemisphere records (e.g. Rasmussen et al., 2006; Broecker, 2010; Fiedel, 2011; Björck, 2013; Carlson, 2013). Cooler temperatures in the Northern Hemisphere led to a decrease in the north-south temperature gradient and subsequently to a southward shift of the mean posi-tion of the ITCZ (e.g. Stouffer et al., 2006; Yancheva et al., 2007; Gasse et al., 2008; Calson, 2013). Distinctly higher precipitation over MSEA is observed already ~11.5 ka BP (unpublished data set), which is earlier than the increase in precipitation over the EAM and IOM sub-systems (Dykoski et al., 2005; Fleitmann et al., 2007; Berkelhammer et al., 2012) (Figs. 11, 17). Wet conditions over MSEA possibly relate to the transport of humidity from Borneo (Partin et al., 2007) over the exposed land areas in the Sundaland and correspond in time with the predomi-nance of La Niña-like conditions at the west coast of Mexico (Marchitto et al., 2010).

The semi-quantitative precipitation reconstructed for MSEA compares moderately well with the qualitative P-E inferred from the new studies at Lake Kumphawapi. The latter studies indi-cate wet conditions between 10 and 7 ka BP and dry conditions after 7 ka BP (Fig. 17).

39

Fig. 18: Reconstruction of the mean position of the ITCZ (dashed red line) in the Asian mon-soon region at 10.8, 9.5 and 6.0 ka BP. The red arrows represent the direction of the ITCZ movement with respect to the reconstructed mean position of the ITCZ. The green lines show the seasonal shift of the present-day ITCZ according to Wang (2009). The grey land areas are the palaeo-shorelines at 10.8, 9.5 and 6.0 ka BP, following Hanebuth et al. (2011).

40

Reconstructed wet conditions over MSEA and the EAM between 10 and 7 ka BP differ from the precipitation pattern over the IOM sub-region, where δ18O values of speleothems from Mawmluh and Quft Cave indicate an insignificant change in precipitation between 10.5 to 9.0 ka BP, but show an increase after 9 ka BP (Fleitman et al., 2005; Berkelhammer et al., 2012). A strengthening of the summer monsoon seems thus to have occurred earlier in the EAM sub-system than in the IOM sub-system. These results support the hypothesis of an asynchronous onset of the Holocene optimum (An, 2000; Herzschuh, 2006; Wang et al., 2010b) (Fig. 17). One possible explanation is the location of MSEA, which acts as a land bridge for the ITCZ.

The dominance of the summer monsoon over the entire Asian monsoon region between 9 and 7 ka BP can be explained by a northward shift in the mean position of the ITCZ, which was triggered by warmer Northern Hemisphere and cooler Southern Hemisphere temperatures as seen in Greenland and Antarctic ice cores, respectively (Linzen and Nigam, 1987; Lindzen and Hou, 1988; NGRIP Members, 2004; Chiang and Bitz, 2005; EPICA Community Members, 2006; Kang et al., 2009; Cvijanovic, 2013) (Fig. 18).

Summer monsoon intensity gradually declined after 7 ka BP following the decrease in inso-lation and the southward shift of the mean position of the ITCZ (e.g. Yang et al., 1995; Bernas-coni et al., 1997; Yan et al., 1999; Mingram et al., 2004; Wang et al., 2010a; Stevenson et al., 2011) (Fig. 18). 5.3 The late Holocene (the last 2 ka): Paper IV Lake Pa Kho was a shallow lake from BC 170 to AD 370, which compares in time with the reappearance of a shallow lake in Lake Kumphawapi about AD 150-350 (Fig. 17). The forma-tion of a peatland suggests that P-E started to decrease ~AD 370. Dry conditions prevailed until AD 800, but were interrupted by wetter conditions between AD 420 and 450. The devel-opment of a wetland ~AD 800 to 970 points to an increase in P-E. A further dry interval may be recognized between AD 970 and 1300, corresponding to the hiatus or to very low peat ac-cumulation. Higher peat accumulation and lower δ13C values indicate an increase in P-E about AD 1300-1450, although overall dry conditions seem to have prevailed until AD 1450. The in-crease in aquatic plant remains, the reappearance of diatoms and the isotope proxies suggest the reestablishment of a wetland and higher P-E after AD 1450.

The hydroclimatic conditions reconstructed for Lake Pa Kho compare moderately well with reconstructed precipitation based on δ18O values of speleothems from Dandak, Jhumar and Wah Shikar Caves in the IOM sub-region, and Wanxiang Cave in the EAM sub-region between BC 170 and AD 800 (Zhang et al., 2008; Berkelhammer et al., 2010) (Fig. 19). The lead and lag of wet or dry intervals can be explained by chronological differences. The reconstructed hy-droclimatic conditions for Pa Kho and for the EAM and IOM regions show a divergent pattern to the precipitation reconstructed from biomarker in Southwest Sulawesi, Indonesia (Tierney et al., 2010). While wet conditions and a stronger summer monsoon prevailed over MSEA and the EAM sub-region between BC 170 and AD 370, precipitation was lower over Southwest Su-lawesi. This would indicate a northward shift in the mean position of the ITCZ (Tierney et al., 2010). Between AD 400 and 800 summer monsoon intensity weakened over EAM and IOM region, but increased over Indonesia, which indicates a southward shift in the mean position of the ITCZ. The ITCZ possibly moved again northward after AD 800, when wet conditions were reconstructed for MSEA and Indonesia, while the IOM and EAM sub-regions remained under dry conditions. The Pa Kho sequence suggests drier conditions and thus a decrease in summer monsoon intensity between AD 1300-1450. This feature may be related to the Walker Circula-tion, because this dry interval is recognized in Indian caves, on Dongdao Island in the South China Sea and in Indonesia (Berkelhammer et al., 2010; Yan et al., 2011; Tierney et al., 2010). Relatively wetter conditions reconstructed for Pa Kho after AD 1450 compare well to records from India, the South China Sea and Vietnam, but diverge from those from Wanxiang Cave in China (Zhang et al., 2008; Berkelhammer et al., 2010; Buckley et al., 2010; Yan et al., 2011).

41

These results suggest that the northward shift in the mean position of the ITCZ may not have reached as far north after AD 1450 as during the former wet intervals.

42

Fig. 19: Reconstructed hydroclimate shifts for Lake Pa Kho compared with selected high-resolution records from the Asian monsoon region for the past 2000 years: (A) δ18O values of speleothem from Wanxiang cave (Zhang et al., 2008); (B) composite δ18O values of speleo-thems from Dandak, Jhumar and Wah Shikar caves (Berkelhammer et al., 2010); (C) δ13C val-ues from Lake Pha Kho (this study); (D) grain size analysis of sediments from Cattle Pond, Dongdao Island (Yan et al., 2011); (E) Palmer Drought Severity Index (PDSI) derived from the Monsoon Asia Drought Atlas (MADA) for the region between 10-20°N and 95-115°E (Buckley et al., 2007, 2010; Sano et al., 2009; Cook et al., 2010; D’Arrigo et al., 2011); (F) δDwax values derived from a marine core from Southwest Sulawesi (Tierney et al., 2010). The dark and light grey bars represent high effective moisture/stronger summer monsoon and low effective moisture/weaker summer monsoon, as interpreted by the original authors. (G) The location of the selected records is shown on the map below (modified from Chawchai et al., submitted).

43

6. Conclusions Mainland Southeast Asia (MSEA) is influenced by the Indian Ocean (IOM) and East Asian monsoon (EAM) sub-systems. It is also the only continental area in the Asian Monsoon region that generates strong convective cells during the initial stage of the summer monsoon onset. MSEA is thus considered a key region for studying the evolution of the Asian monsoon over time.

Published palaeo-records from the Asian monsoon region (including MSEA) have been com-piled, evaluated, compared to climate model simulations, and complemented by new palaeo-proxy data sets derived from lake/peat sequences in northeast Thailand. Thailand is located almost in the middle of MSEA and occupies therefore an important position for tracing past shifts in summer monsoon intensity and for assessing the migration of the mean position of the Intertropical Convergence Zone through time. Palaeo-data - model comparisons suggest a strengthened summer monsoon over MSEA dur-ing the Last Glacial Maximum (LGM) (23-19 ka BP). This can potentially be explained by an increase in land-sea thermal contrast due to the LGM sea level low stand, which exposed the vast Sunda continental shelf. The weakening of the summer monsoon intensity ~20-19 ka BP can be related to the rapid sea level rise ~19.6 ka BP, which led to a reorganization of the at-mospheric circulation over the Pacific Ocean. Reconstructed semi-quantitative precipitation and temperature indicate that climate condi-tions over MSEA between 15 and 5 ka BP alternated between cool/dry and warm/wet states. Precipitation and temperature generally exceeded 1100 mm a-1 and 18 °C, respectively, during this time interval. The semi-quantitative data set derived from MSEA compares well with the Holocene hydroclimatic variability assessed from multi-proxy sediment sequences from Lake Kumphawapi in northeast Thailand. The onset of a stronger summer monsoon at 11.5 ka BP, as reconstructed from MSEA, compares well with the moisture history inferred from palaeo-records in the EAM sub-region, but differs slightly from that of the IOM sub-region, where the onset occurred about 1500 years later. These results support the hypothesis of an asynchro-nous onset of the summer monsoon between the IOM and EAM sub-regions. This asynchrony is possibly related to the land-sea configuration in this part of the Asian monsoon region, where MSEA, which connects the equator and higher latitudes, acted as a land bridge for the north-ward movement of the ITCZ. All data sets from the Asian monsoon region indicate a strength-ened summer monsoon until 7 ka BP, which corresponds to a northward shift in the mean po-sition of the ITCZ. Reconstructed drier conditions in the mid Holocene suggest a weakening of the summer monsoon after 7 ka BP, as a result of the decrease in insolation, and a southward shift in the mean position of the ITCZ. The high-resolution, multi-proxy study of the sediment/peat sequence from Lake Pa Kho in northeast Thailand provides a detailed picture of hydroclimatic changes during the past 2000 years. Intervals of a strengthened summer monsoon are recognized between BC 170 and AD 370, between AD 800 and 960 and since AD 1450 and intervals of a weaker summer monsoon can be observed between AD 370 and 800 and between AD 1300 and 1450. These shifts were mainly controlled by the movement of the ITCZ, except for the drought interval between AD 1300 and 1450, which was possibly triggered by a change in the Walker circulation.

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7. Ongoing works and future perspective 7.1 Asian monsoon variability during the marine isotope stage 3 (MIS 3)

Radiocarbon dating of the 10 m long sediment sequence from Lake Pa Kho shows that the sediments at 6.8 m date to >47 ka BP and that a long hiatus is present between 3.8 and 3.5 m, i.e. between 21.8 ka BP and 2.1 ka BP (BC 170). This means that sediments dating to the LGM to late Holocene are missing in the record. This contrasts with the findings of Penny (2001), who had assumed constant deposition since 40 ka BP.

The fact that sediments dating to between 21.8 and >47 ka BP are present provides an ex-cellent opportunity to study hydroclimatic shifts and Asian monsoon variability during a cli-matically highly variable time interval, namely Marine Isotope Stage (MIS) 3. This time interval (59-29 ka BP) is characterised in ice cores, in marine and lacustrine sediment sequences and in speleothem records by rapid shifts between cold stadials and warm interstadials (Dansgaard et al. 1993; NGRIP Members 2004; Wang et al., 2001; Voelker et al., 2002). Most notable is the asynchrony between Northern and Southern Hemisphere records and the so-called sea-saw effect (Broecker, 2010; Stocker and Johnsen, 2001).

Preliminary geochemical analyses (between 6.8 and 3.8 m depth) of the peaty gyttja in sedi-ment core PK-CP3 from Pa Kho comprise total organic carbon and nitrogen and their isotopes (Fig. 20). Since the sediment composition does not change, I assume that the lake level was relatively stable and that mineral contributions from catchment run off were minor.

Meyers and Teranes (2001) and Meyers (2003) suggested that C/N ratios of <10 and >20 in lake sediment samples would indicate phytoplankton and vascular plant sources, respectively. δ13C values of -30 to -35‰ have been reported for limnic algae and terrestrial C3 plants, while δ13C values of -15 to -10‰ are indicative of C4 plants and values of -11 to -39‰ represent aquatic macrophytes (Meyers and Teranes, 2001; Farquhar et al., 1989).

The TOC values for PK-CP3 range between 35 and 55% between 6.8 and 3.8 m depth (corre-sponding to 47.2-22 ka BP) and do not display any significant changes. Therefore I used the C/N ratio to preliminary subdivide the geochemical record into eight units (Fig. 20). In the bottom unit 1 (47 ka BP), C/N ratios are ~25-29 and δ13C values about -25 ‰, which would suggest a dominance of C3 terrestrial organic matter sources. In unit 2 (47-44 ka BP), C/N ra-tios decline to 21-25 and δ13C values show a slight overall increase to -23 to -19‰, which indi-cates that more aquatic organic matter became incorporated in the sediments. Unit 3 (44-42 ka BP) displays an increase in C/N ratios (22-28) and a further increase in δ13C values to a maxi-mum of -17‰, which could possibly be explained by higher contributions of C4 plants. The C/N ratio is ~25 in unit 4 (42-39 ka BP) and δ13C values decrease to about -21 to -18‰, which suggests terrestrial organic matter sources and possibly higher contributions of C3 plants, as compared to unit 3. The overlying unit 5 (39-35 ka BP) is characterised by an increase in the C/N ratio to 25-30 and a gradual increase in δ13C values from -19 to -17‰. This signifies ter-restrial organic matter input, and an increase in the contribution of C4 plants. The C/N ratio fluctuates ~24 in unit 6 (35-28 ka BP) and δ13C values display an overall increase to -16‰, which points to higher supply of aquatic organic material. In unit 7 (28-25.5 ka BP), C/N ratios increase from 23 to 28 and δ13C values decrease from -19 to -17‰. This would suggest an in-crease in the contribution of C3 plants. The uppermost unit 8 (25.5-22 ka BP) shows a further increase in the C/N ratio and increasing values of δ13C from -17 to -16‰, which points to higher input of C4 plants.

The tentative interpretation of the geochemical record for Pa Kho thus suggests distinct al-ternations in the supply of organic matter to the lake, i.e., dominantly C3 plants, a mix of aquatic and terrestrial organic material, and dominantly C4 plants. The shifts between C3 and C4 vegetation and changes in the supply of organic material into the basin potentially indicate hydroclimatic changes. This would mean wetter conditions (47-44 ka BP), drier conditions (44-42 ka BP), wetter conditions (42-39 ka BP), drier conditions (39-35 ka BP), wetter condi-tions (35-25 ka BP) and drier conditions (25-22 ka BP).

45

Since the age model for PK-CP3 is still in progress, detailed correlations to the Greenland and Antarctic ice cores are premature. However it is obvious that hydroclimate fluctuations are recorded in the sequence, and that these resemble the temperature evolution over the North-ern or Southern Hemisphere.

Fig 20: Results of C/N, δ13C, and TOC measurements in the peaty gyttja of core PK-CP3 from Lake Pa Kho. The sequence between 6.8 and 3.8 m dates to between 47.5 and 21 ka BP. The location of the site is shown in Figure 8. 7.2. Myanmar Project Sediment cores from two crater lakes in northwest Myanmar were collected in November 2013 using a modified Russian corer (7.5 cm diameter, 1 m length) (yellow square) and a grav-ity corer (10 cm diameter, 1 m length) (red square) (Figs. 5, 21). To achieve a complete se-quence, sediment cores were taken with 50-cm overlap with the Russian corer. The gravity cores were sub-sampled in the field in 1.0 and 0.5 cm segments. The sediment cores and sam-ples are stored in the cold room of the Department of Geological Sciences. Preliminary lithostratigraphic descriptions and loss-on-ignition measurements (LOI; 550°C and 950°C for organic matter and carbonate content) have recently been completed. Further analyses (14C chronology, geochemistry) are planned for fall 2014/spring 2015.

46

Fig. 21: Topography of the area hosting the northeast-southwest chain of crater lakes in north-west Myanmar. Sediment cores were obtained from four lakes. The red squares mark the loca-tion of gravity cores and the yellow and red squares the locations where both long and short sequences were obtained.

47

8. Acknowledgements First of all, I express my deep gratitude to Barbara Wohlfarth, my main supervisor, for a

challenging and interesting research topic and her encouragement and support to me through-out this project.

I would like to thank Jenny Brandefelt, my co-supervisor, for the climate model data output, valuable advice and comments on Paper I. The completion of this study could not have been accomplished without the support of my colleagues, Sakonvan Chawchai, Ludvig Löwermark, Sheri Fritz, Wichuratree Klubseang, Suda Inthongkeaw, Kweku Afrifa Yamoah, Minna Väliranta, Heike Siegmund, Malin Kylander, Carl-Magnus Mörth, Maarten Blaauw, Paula Reimer and Rienk Smittenberg.

I would also like to express my gratitude to Hildred Crill for language advice, Klara Hajnal and Carina Johansson for their help in the laboratory, Francis Freire for GIS advice, Agatha de Boer, Daniele Reghellin, Francesco Muschitiello, Frederik Schenk and Robert Graham for help-ful discussions, and Oscar Fransner, Ludvig Löwermark, and Kajsa Markdahl for the Swedish abstract.

Finally I would like to thank Barbara Kleine, Daniele Reghellin, Duc Nguyen, Francis Freire, Jose Godinho, Kweku Afrifa Yamoah, Oscar Fransner, Sakonvan Chawchai, Stefano Bonaglia and Zhang Xiao-Jing for our good time together in Stockholm.

Swedish Research Council (VR) grants 621-2008-2855, 348-2008-6071 and 621-2011-4684 provided financial support for this research.

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Appendix I Palaeo-records and palaeo-proxies used for the compilation in Paper I. The age assignments are based on published 14C, TL and U/Th dates. The number of dates between 25-17 ka BP for each sequence is given in brackets.

Site

no. Site Name

Lat

(°)

Long

(°)

Eleva-

tion (m a.s.l.)

Archive Proxy Age assign-

ment References

1 Kosipe valley -8.45 147.20 1965 T Pollen 14C (1) Hope (2009) 2 Situ Bayong-

bong swamp -7.18 107.28 1300 T Pollen 14C (1) Stuijts et al.

(1988)

3 Bandung basin

-7.00 108.00 665 T Pollen 14C (1) van der Kaars and Dam (1997)

4 BAR94-42 -6.75 102.42 -2542 M Pollen 14C (2) van der Kaars et

al. (2010) 5 SHI9014 -5.77 126.97 -3163 M Pollen 14C (1) van der Kaars et

al. (2000)

6 Sentarum lake

0.73 112.10 35-50 T Pollen 14C (1) Anshari et al. (2001)

7 di Atas lake -1.07 100.77 1535 T Pollen 14C (2) Newsome and

Flenley (1988) 8 Pea Sim-sim

swamp

2.29 98.89 1450 T Pollen 14C (5) Maloney (1980)

9 Pee Bullok swamp

2.28 98.98 1400 T Pollen 14C (4) Maloney and McCormac (1996)

10 K-12 2.69 127.74 -3510 M Pollen 14C (1) Barmawidjaja et

al. (1993) 11 Tasek Bera

basin 3.06 102.64 20-30 T Hard-

wood remain

14C (1) Wüst and Bustin (2004)

12 Cave in Gunung Buda National Park

4.00 114.00 ~1000 T δ18O U/Th (9) Partin et al. (2007)

13 SO18302 4.15 108.57 83 M Pollen 14C (1) Wang et al. (2009)

14 SO18300 4.35 108.65 91 M Pollen 14C (1) Wang et al. (2009)

15 Core 17940 20.12 117.38 -1727 M Pollen 14C (2) Sun et al. (2000) 16 Tianyang

basin 20.78 110.03 120 T Pollen 14C (2) Zheng and Lei

(1999) 17 Huguang lake 21.15 110.28 23 T Pollen 14C (2) Wang et al.

(2010b) 18 Toushe Basin 23.82 120.88 650 T Pollen 14C (3) Liew et al.

(2006) 19 DGKS9603 28.15 127.27 -1100 M Pollen 14C (2) Xu et al. (2010) 20 Jintanwan

Cave 29.48 109.53 460 T δ18O U/Th (3) Cosford et al.

(2010)

ii

21 Hulu Cave 32.30 119.17 100 T δ18O U/Th (3) Wang et al. (2001)

22 Songjia Cave 32.41 107.41 ~680 T δ18O U/Th (2) Zhou et al. (2008)

23 Weinan sec-tion

34.40 109.50 600-1100

T Pollen 14C (1) Sun et al. (1997)

24 Beizhuang-cun section

34.33 109.48 600-1100

T

Pollen

14C (2)

Wang and Sun (1994)

25 Pyonggeo-dong archae- ological site,

35.17 128.06 100-300 T Pollen 14C (3) Chung et al. (2006)

26 Biwa lake 35.25 136.05 85 T Pollen 14C (2) Hayashi et al.

(2010) 27 A paddy field

at Iwaya, Fukui prefec-ture

35.52 135.88 20 T Pollen 14C (1) Takahara and Takeoka (1992)

28 Mikata Lake 35.56 135.89 0 T Pollen 14C (2) Yasuda (1982) 29 SK-157-14 5.18 75.91 -3306 M δ18O 14C (1) Ahmad et al.

(2008) 30 Horton plains 6.81 80.83 2100-

2300

T Pollen 14C (2) Premathilake

(2006) Premathilake and Risberg (2003)

31 Nilgiri hills 11.25 76.67 2200 T δ13C 14C (1) Rajagopalan et

al. (1997) 32 RC12-344 12.46 96.04 -2140 M δ18O 14C (3) Rashid et al.

(2007) 33 Moomi cave, 12.50 54.00 ~1000 T δ18O U/Th (8) Shakun et al.

(2007) 34 SO90-137KA 23.12 66.48 -573 M δ18O 14C (4) van Rad et al.

(1999) 35 Bharatpur

Bird Sanctu-ary wetland

27.12 77.52 174 T Pollen 14C (1) Sharma & Chat-terjee (2007)

36 Phulara pa-laeolake

29.33 80.13 1500-1700

T Pollen 14C (2) Kotilia et al. (2010)

37 Kathmandu basin

27.67 85.22 1303 T Pollen 14C (3) Fujii and Sakai (2002)

38 Shudu lake, 27.90 99.95 3630 T Pollen 14C (4) Cook et al.

(2011) 39 Ren Co 30.73 96.68 4450 T Pollen 14C (3) Tang et al.

(2000) 40 Tham Rod

archaeolo-

gical site

19.57 98.89 600-1170

T Fauna remains

TL (1) Wattanapituksa-kul (2006)

iii

Appendix II Palaeo-vegetation records used in Paper II.

Site no.

Site name Lat (°)

Lon (°)

Eleva-tion

(m a.s.l.)

No. of 14C

samples

Calibrated age interval

(ka BP)

References

1 Red River core VN (RR_VN)

20.41 106.38 0.6 3 8.39-10.13 Li et al. (2006)

2 Red River core GA (RR_GA)

20.25 106.52 1.1 1 10.46-10.68 Li et al. (2006)

3 Kumphawapi (KMP) 17.18 103.03 170 5 6.21-14.55 Kealhofer and Penny (1998)

4 Pa Kho (PK) 17.10 102.93 ~170 2 7.48-12.87 Penny (2001)

5 Nakhon Nayok (NYY) 14.00 100.92 ~2.0 2 6.84-8.97 Punwong (2007)

6 Senanivate (SNN) 13.85 100.60 2.0 3 5.37-8.64 Somboon (1988)

7 Kara Lake (KRL) 13.55 107.13 150.0 4 5.31-10.60 Maxwell (2001)

8 Tonle Sap (TLN) 12.59 104.29 5.0 3 5.64-7.98 Penny (2006)

9 Mekong River delta core PK (MK_PK)

11.19 105.28 6.0 7 6.27-8.12 Li et al. (2012); Tamura et al. (2009)

10 Mekong River delta core PSG (MK_PSG)

10.96 105.06 3.5 14 5.76-9.07 Li et al. (2012); Tamura et al. (2009)

11 Nong Thale Song Hong (NTSH)

7.87 99.48 ~100 2 7.16-12.87 Maloney (1999)

12 Thale Noi (TLN) 7.20 100.47 ~4.0 2 7.37-8.11 Horton et al.

(2005) 13 Songkhla Lake (SKL) 7.27 100.50 8.0 1 6.74-7.03 Rugmai (2006)

14 Tasek Bera (TSB) 3.81 102.40 25.0 1 5.05-5.30 Morley (1981); Maloney (1985)

15 Nee Soon (NS) 1.40 103.80 <10.0 4 5.99-13.79 Taylor et al. (2001)

16 Kwan Phayao (KPY) 19.17 99.87 380 1 11..90-11.36 Penny and Keal-hofer (2005)