Biogeochemical evidence for environmental and vegetation changes in peatlands from the middle Yangtze river catchment during the medieval warm period and little ice Age

Abstract

The reconstruction of past climates and vegetation can provide valuable information for understanding the mechanisms of the variability of the East Asian summer monsoon in eastern China. In this study, organic geochemical evidence from the compositions of sedimentary leaf-wax stable isotopes and n-alkane parameters investigates the changes in vegetation and climate over the last 1200 years in the Xiyaohu peatland, Jiangxi Province, southeast China. Combined with temperature records, three climatic periods are presented: (a) a warm and humid period with an increase in C₄ plants from 900 to 1450 AD, which coincides with the Medieval Warm Period (MWP); (b) a cool and dry period with the expansion of C₃ plants from 1450 to 1800 AD, coinciding with the Little Ice Age (LIA); and (c) the Present Warm Period (PWP) from 1800 AD until the present, with warm and wet conditions. The sub-stages within the MWP and LIA intervals are also presented. The earlier MWP stage (900–1125 AD) was drier than the latter one (1125–1450 AD), and the earlier LIA stage (1450–1650 AD) was drier than the late LIA (1650–1800 AD). Increased solar irradiance and enhanced El Niño activities are related to the warm and humid climate during the MWP and PWP, whereas reduced solar irradiance and La Niña activities correspond to the cool and dry climate during the LIA. The present results provide insights into paleoclimatic changes in eastern monsoonal China and provide an understanding of centennial-scale climatic fluctuations and their driving factors.

Keywords

δD LIA MWP

Introduction

The Asian monsoon is a key component of the global climate system. Summer monsoon-derived rainfall is crucial to terrestrial ecosystems and human societies in East Asia (An, 2000). The last 1200 years include the Medieval Warm Period (MWP) and Little Ice Age (LIA), which are multi-centennial-scale global climatic events (Mann et al., 1999). The MWP prevailed in various parts of the world, and historical documents and paleoclimate reconstructions show that the temperature increased during the MWP. Continental and oceanic temperature reconstruction records demonstrate that the LIA is the nearest mountain-glacier expansion episode on a centennial time scale and featured colder conditions than those of medieval times (D’Arrigo et al., 2006; Holzhauser et al., 2005; Kayastha and Harrison, 2008; Lorrey et al., 2014; Matthews and Briffa, 2016). Paleoclimate reconstructions and simulations show that precipitation during the MWP and LIA presents different patterns in China. The wet–dry pattern of the eastern monsoon region of China is unstable and controversial (Chen et al., 2015). The middle Yangtze River in the East Asia monsoon area is typically recognized as one of the origins of Chinese civilization and supports prosperous rice cultivation, which depends primarily on water availability and mild temperatures (Normile, 1997). Climatic variability during the MWP and LIA profoundly affects southeast China (Zhao and Piperno, 2000), therefore, a clear understanding of past monsoonal climate changes and their drivers, particularly during key periods such as the MWP and LIA, is essential to study the response of terrestrial ecosystems to climate change and its influence on human development.

n-alkanes are biomarkers that are composed of a chain of methylene groups and epidermal cells by the elongation of C₁₆ fatty acids. In the absence of functional groups, they are thermally stable and relatively resistant to degradation (Naafs et al., 2019). n-alkane distribution and stable isotopic composition (δD and δ¹³C) have been widely used to reconstruct various paleoenvironmental changes, such as temperature variability, paleo-vegetation evolution, and hydrological changes (Feakins et al., 2019; Ficken et al., 2000; Wang et al., 2017).

Ombrotrophic peatlands ideal archives for high-resolution paleoclimate studies owing to their high accumulation rates and sensitivity to climate change (Gałka et al., 2017; Schittek et al., 2016). Studies have been conducted in South China using lipid biomarker approaches (Huang et al., 2018; Zheng et al., 2009; Zhou et al., 2005). The hydrogen isotope ratios of n-alkanes from the Poyang Lake sequence were used to reconstruct precipitation changes (Yao et al., 2015). n-alkane distributions and other proxies from a peat deposit from southern China were applied to investigate paleoclimate changes from 18 ka (Zhou et al., 2005). However, high-resolution studies of n-alkane δD and δ¹³C values in peat from South China are relatively few.

Xiyaohu (XYH) is a relatively elevated peatland in Jiangxi Province in the south of the middle Yangtze River catchment, and multiple proxies from the peatland, including pollen, phytolith, and biomarkers have been used to reconstruct vegetation changes (Cui et al., 2018; Zhang et al., 2019; Pang et al., under review). A previous study conducted a grain-size analysis to study climate change in the past 2000 years at Xiyaohu (Li et al., 2017). In this study, sediment leaf-wax n-alkane distributions, δD, and δ¹³C, from the XYH peat sequence for the last 1200 years were analyzed. Leaf-wax δD records δD precipitation in the middle Yangtze River, and δ¹³C typically indicates photosynthetic pathways for C₃ and C₄ plants (Collister et al., 1994). This study aimed to detect the changes in the botanical sources of the peatland and to reconstruct paleoclimate changes during the past 1200 years. Considering the integration of human activities and climatic changes, the results were compared with previously published proxies and historical documents to analyze different anthropogenic and natural influencing factors.

Study area

The XYH peatland (28.44°N, 115.40°E; Figure 1b) is located on the northwest main peak of Mount Xishan in Jiangxi province, in the middle reaches of the Yangtze River. The bottom of the peatland is flat and covered by a silty clay layer and heavily weathered exposures of granite bedrock (Figure 1c). Rainfall is the primary water source, and hygrophilous vegetation is abundant because of recharge from slope runoff. This provides an ideal situation for the formation of ombrotrophic peatland. The XYH peatland is the largest and thickest at Mount Xishan, with a surface area of 12,000 m² and a maximum depth of 3.8 m.

Figure 1.

(a) Monthly temperature and precipitation from the nearby Nanchang meteorological station. (b) The Location of XYH (red star) and another site discussed in the study (yellow star). The dashed line represents the boundary of the modern summer monsoon, and the arrows represent two storm trajectories (Chen et al., 2010). The pink shadow represents the Jiang-Nan area (Zheng et al., 2006). The base map is provided from the ArcGIS Hub. (c) The General cross-sectional view of the studied peatland in Xishan Mountain (modified from Zhang et al., 2019).

The study area is influenced by the humid subtropical East Asian summer monsoon and experiences summer-dominated precipitation, with maximum precipitation from April to June (Figure 1a). The total annual precipitation is 1400 mm, and the annual mean temperature is 17°C–18°C with a short spring and autumn but a long summer and winter (Bureau, 2015).

The vegetation on Mount Xishan changes with elevation. The foothills contain economic crops, namely two-season croplands, rapeseed, and tea-oil tree forests. The hillsides are dominated by widespread scrubs (i.e. Loropetalum chinense, Caccinium bracteatum, and Rhododendron simsii scrub), grass-forb communities (Arundinella setosa and Miscanthus sinensis), and sporadic forests (Phyllostachys pubescens, Pinus massoniana, Loropetalum chinense, Rhododendron simsii, and Cunninghamia lanceolate). Areas with elevation above 700 m feature the grass group, comprising swamps (Miscanthus sacchariflorus, Phragmites communis, Carex spp., and Juncus effuses), meadows (Carex cinerascens), and scrub (Salix spp. and Lespedeza bicolor). Vegetation in the peatland area includes widespread grasses, sporadic scrubs, meadows, and swamps) (Vegetation Atlas of China, 2001). The dominant vegetation in the valley includes Scirpus cyperius, Salix brachypoda, and Sphagnum assemblages.

Materials and methods

Core sampling and stratigraphy

An XYH sediment core was drilled from the center of the XYH peatland using a manual sampling peat drill with a PVC tube with a length of 300 cm and an inner diameter of 60 mm. The total depth of the sample (300 cm) was divided into three units based on lithology. The top (0–120 cm) of the core consisted gray-black clay with abundant plant residue. From 120 to 215 cm, gray–yellow and dark gray muddy sand is the primary component. The bottom (215–300 cm) changed from taupe to gray-black peat soil (Figure 2).

Figure 2.

Lithology of the XYH peatland and age-depth model. Gray dots indicate the 95% probability intervals. Average linear sedimentation rates (sr) are also shown.

Core chronology

All procedures were conducted in the Pollen and Paleoecology Lab, Department of Geography and Ocean Science, Nanjing University. Radiocarbon dating of sporopollen concentrates, plant residues, and charcoal samples (total of 12) was conducted by acceleration mass spectrometry (AMS) at the Beta Analytic Laboratory (Florida USA; Table 1). The Bayesian software Bacon “Winbacon2.2” (Blaauw and Christeny, 2011) and R statistical software were used to develop an age model using 12 AMS radiocarbon dates (Blaauw and Christeny, 2011). CALIB Rev. 5.0.2 was used to calibrate the radiocarbon age to calendar years. Generally, the section from 86 to 120 cm displayed notably lower sedimentation rates than the one from 153 to 300 cm. The top section from 0 to 86 cm has a medium sedimentation rate.

Table 1.

AMS ¹⁴C dating and correct results of XYH core.

Lab No.	Sample ID	Depth (cm)	Material dated	AMS ¹⁴C date (year BP)	δ¹³C (‰)	Calibrated age (AD)
Beta-465662	XYH14-51	51	Pollen concentrates	110 ± 30	−27.1	1802–1938
Beta-432598	XYH14-86	86	Plant residues	180 ± 30	−12.4	1655–1695
Beta-465664	XYH14-102	102	Pollen concentrates	210 ± 30	−21.2	1734–806
Beta-422837	XYH14-120	120	Pollen concentrates	510 ± 30	−21.9	1400–1440
Beta-432599	XYH14-140	140	Pollen concentrates	540 ± 30	−19.4	1320–1350
Beta-465665	XYH14-153	153	Pollen concentrates	600 ± 30	−18.8	1296–1409
Beta-422838	XYH14-165	165	Plant residues	820 ± 30	−21.2	1165–1265
Beta-465669	XYH14-232	232	Pollen concentrates	1100 ± 30	−20.5	886–1013
Beta-422839	XYH14-245	245	Charcoal	880 ± 30	−28.4	1045–1095
Beta-465670	XYH14-261	261	Pollen concentrates	1080 ± 30	−22.1	938–1018
Beta-465671	XYH14-281	281	Pollen concentrates	1050 ± 30	−21.4	948–1026
Beta-422841	XYH14-300	300	Plant residues	1220 ± 30	−26.3	690–750

Lipid extraction

The core was selected at 5- and 4-cm intervals through the top and bottom 150 cm, respectively; 68 samples (5 g/sample) were freeze-dried and homogenized. Sediments were solvent-extracted with 9:1 v/v dichloromethane (DCM) to methanol (MeOH) at 30°C for 30 min and 1500 psi for 5 min using an accelerated solvent extraction system. Total lipid extracts were separated using column chromatography through silica gels into the hexane and neutral fractions, using hexane and 9:1 v/v DCM: MeOH, respectively. The hexane and neutral fractions contain n-alkanes and the remaining parts of the lipid, such as long-chain ketone, respectively.

n-alkanes analysis

Compound identification and the molecular abundance distribution of long-chain n-alkanes were detected by gas chromatography (GC, Agilent 7890B) equipped with a flame ionization detector (FID) for separation and quantification. The FID temperature was 310°C, the H₂ speed was 30 mL/min, the airflow speed was 400 mL/min, and each sample tested was 1 µL. Individual n-alkane identification was performed by comparing the retention times of n-alkanes in samples with the aforementioned standards. We report the average chain length (ACL) (Poynter et al., 1989), carbon preference index (CPI) (Bray and Evans, 1961; Ortiz et al., 2010), and P_aq (Ficken et al., 2000):

\begin{array}{l} {ACL}_{23 - 33} = (23 \times C_{23} + 25 \times C_{25} + \dots + 33 \times C_{33}) \\ ⁄ (C_{23} + C_{25} + \dots + C_{33}) \end{array}

CPI = \frac{1}{2} \times [\begin{array}{l} (\frac{C_{25} + C_{27} + C_{29} + C_{31} + C_{33}}{C_{24} + C_{26} + C_{28} + C_{30} + C_{32}}) \\ + (\frac{C_{25} + C_{27} + C_{29} + C_{31} + C_{33}}{C_{26} + C_{28} + C_{30} + C_{32} + C_{34}}) \end{array}]

Paq = (C_{23} + C_{25}) ⁄ (C_{23} + C_{25} + C_{29} + C_{31})

Hydrogen and carbon isotope analysis

δD and δ¹³C values of long-chain n-alkanes were measured with a gas chromatographer (Trace GC Ultra) connected to a Thermo Fisher Delta V Advantage isotope ratio mass spectrometer. δD values were normalized to the Vienna Standard Mean Ocean Water (VSMOW) hydrogen isotopic scale. We measured the δ¹³C values of leaf-wax n-alkanes. Analytical conditions were the same as the abovementioned, and the δ¹³C values were normalized to the Vienna Pee Dee Belemnite (VPDB) carbon isotopic scale.

Results

n-alkane distributions

The carbon numbers in the peat samples ranged from C₁₇ to C₃₅. The relative abundances of n-alkanes in the peat samples show the dominance of mid- to long-chain n-alkanes with variable major peaks. Figure 3b and c show two distinct n-alkane distribution patterns. Pattern A (a single-peak type) features the highest peak at C₂₉ or C₃₁, and pattern B (a double-peak type) features the highest peak at C₂₅, C₂₇, C₂₉, or C₃₁. Pattern B has considerably more mid-chain n-alkanes than pattern A.

Figure 3.

Results of n-alkanes and its parameters values. (a) n-alkane distributions and P_aq, ACL, and CPI values of the XYH core. (b) and (c) n-alkane distribution patterns. P-A and P-B in (a) are the acronyms of pattern A and pattern B, respectively.

Figure 3a shows the variations in the n-alkane distributions and parameters; ACL, P_aq, and CPI values. The ACL values generally fluctuated from 28.0 to 30.7 with a mean of 29.0, from 813 to 2010 AD. The values remained low from 813 to 1025 AD then rapidly increased to a mean of 29.5 at the end of 1125 AD. The mean ACL value fell to 28.9 from 1125 to 1450 AD before growing significantly and remaining stable for more than 200 years. From 1650 to 1925 AD, ACL values reached a minimum of 28.4 and then recovered slightly to 28.9. P_aq values ranged from 0.11 to 0.40 with a mean of 0.24. The average P_aq values had a relatively low mean of 0.19 from 813 to 1025 AD, increased to 0.26, from 1125 to 1440 AD, varied for a decade, and increased to a peak of approximately 0.32 between 1650 and 1925 AD. CPI values ranged from 1.24 to 6.01 with a mean of 2.99. The values remained high from 813 to 900 AD with a mean of 3.48, and then sharply decreased to 1.36 at approximately 1000 AD. After a rapid increase from 1025 to 1125 AD, the values decreased, with a mean of 2.01 from 1125 to 1250 AD and a mean of 2.26 from 1250 to 1450 AD. The CPI values then increased considerably from 1450 to 1800 AD with a mean of 3.68, followed by a decrease in CPI from 1800 to 1925 AD.

Hydrogen and carbon isotope compositions of n-alkanes

The δD values of C₂₇, C₂₉, and C₃₁ n-alkanes (Figure 4a) ranged from −178.00‰ to −199.00‰, −179.00‰ to −200.95‰, and −182.00‰ to −233.00‰, respectively. The C₂₇ n-alkanes exhibited less covariance with C₂₉ (r² = 0.2809, p = 0.011) and C₃₁ n-alkanes (r² = 0.175, p = 0.0124), whereas C₂₉ and C₃₁ n-alkanes exhibited similar covariances (r² = 0.5225, p < 0.0001). The δD_C29 values decreased from 813 to 1125 AD, increased by 9‰ from 1125 to 1300 AD and 1300 to 1650 AD, and thereafter decreased significantly from −189‰ to −199‰ and fluctuated at −199‰ for approximately one decade. The δD_C29 values increased again and fluctuated at −195‰ after 1650 AD. The weighted mean values of δD_C29 and δD_C31 were calculated to minimize the influence from vegetation and ecosystem type variation. The δ¹³C values of C₂₇, C₂₉, and C₃₁ n-alkanes (Figure 4b) ranged from −24.9‰ to −33.2‰, −25.9‰ to −33.3‰, and −26.9‰ to −33.9‰, respectively, and exhibited significant covariation over time. The weighted average δ¹³C_mean values of C₂₇, C₂₉, and C₃₁ n-alkane δ¹³C values, which ranged from −26.0‰ to −33.5‰ with a mean value of −30.3‰, were presented. The lowest δ¹³C_mean values (−33.47‰) were reported at approximately 820 and 1700–1800 AD, and the largest δ¹³C_mean values (−26.01‰) occurred at approximately 1400 AD. A growth of 4.4‰ occurred from 813 to 1025 AD, followed by a rapid decrease. The δ¹³C_mean values gradually increased by 2.2‰ at 1440 AD and generally decreased by 4.28‰ from 1440 to 1650 AD, with several short intervals of increase, before slowly growing and stabilizing at approximately −32.9‰.

Figure 4.

Results of hydrogen and carbon isotope composition of n-alkanes. (a) δD values of C₂₇, C₂₉, and C₃₁ n-alkanes and δD_mean values. (b) δ¹³C values of C₂₇, C₂₉, and C₃₁ n-alkanes, and δ¹³C_mean values.

Discussion

Origin of sedimentary n-alkanes and proxies interpretation

Modern studies have illustrated that short-chain n-alkanes with 15–19 carbon atoms are primarily produced by algae and photosynthetic bacteria. Moss and aquatic macrophytes, including submerged, floating, and emerged plants, biosynthesize multiple mid-chain n-alkanes with 21–27 carbon atoms. Terrestrial plants predominantly produce odd-numbered long-chain n-alkanes with 25–37 carbon atoms (Barnes and Barnes, 1978; Cranwell, 1984; Cranwell et al., 1987; Eglinton and Hamilton, 1967). In this study, the proportions of mid- to- long-chain n-alkanes (C₂₁–C₃₁) accounted for approximately 95% of the total. Although the proportional n-alkanes input from diverse plants does not accurately indicate these amounts, terrestrial plants, moss, and aquatic plants with little input of microorganisms or diatoms are considered the dominant contributors to sedimentary n-alkanes. The present results indicate a prominent change in the proportion of aquatic and terrestrial plant inputs. P_aq was defined to reconstruct the input of aquatic macrophytes in peat (Ficken et al., 2000). ACL_23–33 values are the indicators of temperature, precipitation, and effective moisture depending on the dominant regional climatic control factors (Naafs et al., 2019). In this study, ACL_23–33 values exhibited an opposite trend to that of P_aq change, indicating that hydrological changes caused a decrease in aquatic macrophytes and an increase in terrestrial plants. The two n-alkane distribution patterns alternated through the sedimentary core (Figure 3a). During the time intervals of 900–1125 AD, 1450–1650 AD, and 1800–2000 AD, C₂₉ and C₃₁ n-alkanes were the most abundant homologues and exhibited pattern A one-peak mode. From 813 to 900 AD, 1125 to 1450 AD, and 1650 to 1800 AD, pattern B that represented two-peak variants had higher mid-chain n-alkanes C₂₅ than pattern A. Pattern A represents the period in which terrestrial plants thrived and provided the principal sediment contribution. However, pattern B indicates the combined input of aquatic macrophytes and terrestrial plants. This inference is consistent with the variation in P_aq. When n-alkanes were distributed as pattern A, the P_aq values and C₂₅ n-alkane concentrations were remarkably low, that is, from 1450 to 1650 AD. Pollen records in the same study core exhibited a conspicuous decline in wetland herbs and increase in Poaceae. When n-alkanes were distributed as pattern B, the P_aq values were higher as the C₂₉ and C₃₁ n-alkane concentrations increased. During these periods, the abundance of wetland-herb pollen increased significantly, whereas that of Poaceae, birch, and alder pollen decreased. Therefore, the primary n-alkane sources were the peatland and surrounding hills, as recorded by regional paleoclimate changes (Cui et al., 2018).

The δ¹³C values of n-alkanes from sedimentary leaf wax lipids are typically used as the indicators of the relative contribution of the different photosynthetic pathways, C₃ and C₄. Almost all trees, cool-season grasses, and cool-season sedges use the C₃ photosynthetic pathway, whereas the C₄ photosynthetic pathway is used by warm-season grasses and sedges (Castañeda et al., 2009). Submerged and surface aquatic plants typically use the carbon in CO₂ to maintain regular photosynthesis. C₃ and C₄ plants have different CO₂ fixation modes, whereby, C₃ plants incorporate CO₂ with the Calvin cycle and show enriched ¹²C (depleted δ¹³C values range from −34‰ to −24‰). C₄ plants fix CO₂ with the Hatch–Slack cycle and exhibit enriched ¹³C (enriched δ¹³C values ranging from −19‰ to −6‰) (Lichtfouse et al., 1994). The δ¹³C values from the Beta lab showed a significant difference from the one from our n-alkanes. We noticed that most of the AMS δ¹³C values ranged from −18‰ to −22‰, but some of them have an obvious deviation (samples at the depth of 51, 86, 245, and 300 cm). According to Beta Analytic White Paper (www.radiocarbon.com/accelerator-mass-spectrometry.htm), the AMS δ¹³C values were measured on accelerator mass spectrometry with a series of pretreatments such as the combustion of the samples, the CO₂ graphitization reaction, and the AMS deflection during the testing. It is widely accepted that isotopic fractionation during the processes occurs, which could lead to 0‰–20‰ anomaly from the real δ¹³C from the samples tested on gas chromatographer in our lab. In addition, the AMS δ¹³C values could be used to ¹⁴C isotopic correction and calculation, and cannot represent the condition of the original samples. The differences that occurred from the table might come from the different materials’ pretreatment and the AMS testing processes.

Relative humidity and vegetation type influence leaf-wax δD. δD values in vascular plant leaf waxes, such as C₂₉ n-alkanes, have been used to identify the δD_{precipitation} source water (in the present study area, the source water is primarily precipitation) values during growth periods (Sachse et al., 2006). Nevertheless, other factors, such as leaf-water evaporation and species distribution, can affect leaf-wax δD. Previous studies have shown that ε_wax/p (the apparent fractionation between the leaf-wax n-alkane δD and precipitation δD_{precipitation}) is affected by multiple factors, such as the growing season length, plant life form, precipitation amount, and evapotranspiration (Sachse et al., 2012). However, ombrotrophic peatlands are singular in vegetation and environment in some ways, minimizing the influence of leaf-water evaporation and species (Nichols et al., 2009). Leaf water evaporation is primarily influenced by the leaf-to-air vapor difference, which depends on relative humidity and stomatal conductance (Hou et al., 2008). The relative humidity in ombrotrophic peatlands fluctuates between 75% and 85%, and therefore, the effect of variations in relative humidity on the δD of leaf wax in the peatland could be minimized. A previous study showed that the impact of relative humidity on plant leaf-wax δD was rather small in growth experiments, with only a 5.7‰ deviation for a relative humidity difference of 50% (Hou et al., 2008). The stratigraphic pollen record shows that the primary assemblages in the area are Pinus, Castanopsis, Poaceae, Artemisia, Cyperaceae, and Apiaceae (Cui et al., 2018); 40% of the pollen assemblage was dominated by upland herbs. Although regional vegetation changes substantially throughout the period, the site remains a peatland with relatively low diversity. Most of the vegetation changes in peatland are between vascular plants and the Sphagnum, and the sediment leaf wax originates from peatland plants in situ. The differences between species can be minimized. Herein, the δD of vascular plants was used to interpret the δD_{precipitation}. The δD of stored peatland water represents the mean annual precipitation and is more negative with abundant precipitation.

Aquatic plants are dominated by mid-chain n-alkanes, whereas terrestrial plants are dominated by long-chain n-alkanes. The C₂₃/C₃₁ n-alkane ratio was used to reconstruct the input of Sphagnum in modern peat (Nott et al., 2000). However, the C₂₃/C₃₁ n-alkane ratio can also be affected by environmental and biosynthetic processes. The P_aq index indicates the relative contributions of aquatic macrophytes to peatlands and the changes in peat hydrology (Nichols et al., 2006; Routh et al., 2014; Zhou et al., 2005). High P_aq values (>0.7) indicate a significant contribution of aquatic macrophytes, whereas low P_aq values (<0.1) indicate a contribution from terrestrial plants. However, P_aq values cannot be applied to distinguish the similar lipid distributions of Sphagnum and aquatic macrophytes (Baas et al., 2000; Ficken et al., 2000), particularly when plant macrofossils are absent or poorly preserved (Ortiz et al., 2010). In this study, the P_aq values complement the C₂₃/C₃₁ n-alkane ratio to reflect the Sphagnum (aquatic macrophyte) input relative to the terrestrial plants in peats. The increase in C₂₃/C₃₁ n-alkane ratio and P_aq indicates a large amount of aquatic macrophytes, which typically flourish under wet and humid conditions, and vice versa.

The CPI values of n-alkanes indicate the predominance of odd-numbered over even-numbered n-alkanes. Although CPI from fresh plant is high (e.g. 4–40), it decreases with diagenesis and microbial input, reaching 1. The mean CPI value in this study was 2.99, which is most likely caused by the bacterial diagenesis of n-alkanes and potential production from bacterial long-chain n-alkanes. This biosynthesis process often occurs at elevated temperatures. In this circumstance, CPI values cannot be considered as a direct proxy to infer the changes in paleoclimate but can support other temperature records in the nearby area.

Reconstruction of the paleoclimate and vegetation

Based on the n-alkanes and stable isotope records from the XYH peat core, we divided the climate evolution into five stages (Figure 5).

Figure 5.

Comparison of different climatic records. (a) P_aq record (this study), (b) C₂₃/C₃₁ n-alkane ratio record (this study). (c) δD_mean records (this study). (d) Wetland-herb pollen concentration of core XYH (Cui et al., 2018). (e) Arid index Iph phytolith records from XYH (Zhang et al., 2019). (f) Speleothem δ¹⁸O record from Cave Heshang (Hu et al., 2008). (g) Dry-Wet index in Jiang-Nan area, approximately 25°N–31°N (Jingyun et al., 2006). (h) δD_C31 record from Poyang Lake (Yao et al., 2015). (i) Warm index Iw phytolith records from XYH (Zhang et al., 2019). (j) δ¹³C_mean record (this study). (k) Mean air temperature anomaly in East China over the past 3000 years (Liu, 2004 after Zhang et al., 2000).

Stage I (900–1125 AD) (red bar in Figures 5 and 6): The sediment was primarily dark peat with approximately 10 cm of silty mud. The sedimentation rates increased from 0.27 to 0.37 cm/year. The low P_aq and C₂₃/C₃₁ ratios, compared to the later period, revealed a smaller amount of aquatic plant input, which was correlated with relatively less negative δD_mean values. The overall decrease in δD_mean during this period implied a wetting trend. The overall dry conditions weakened the growth of aquatic plants, which is supported by the relatively low pollen concentration of wetland herbs (Cui et al., 2018) and phytolith index (Iph) values (Zhang et al., 2019). The δ¹³C_mean values were relatively negative but increased, as C₃ biomass was the dominant and C₄ plants increased. Studies have shown significantly negative correlations between δ¹³C values and mean annual precipitation for C₃ plants and a significantly positive correlation between δ¹³C values and mean annual temperature for C₄ plants (Liu and An, 2020). From the compiled sedimentary and historical records of temperature in east China (Liu, 2004; Zhang et al., 2000), the temperature increased significantly compared to that from 813 to 900 AD. This increase in temperature corresponded with the high phytolith warm index Iw values and likely implied warming conditions. The δ¹³C_mean values increased as the number of C₄ plants increased, and therefore, it is suggested that higher temperatures benefit C₄ plant growth.

Figure 6.

(a) P_aq record (this study). (b) C₂₃/C₃₁ ratio record (this study). (c) δD_mean record (this study). (d) δ¹³C_mean record (this study). (e) El Niño 3.4 index (Emile-Geay et al., 2013). (f) Solar irradiance record (Bard et al., 2000). (g) Sea surface temperature (SST) of the eastern equatorial Pacific (Conroy et al., 2010). (h) Historic population change records from the historical documents in the Library of Nanjing University.

Stage II (1125–1450 AD) (yellow bar in Figures 5 and 6): The sediment was primarily silty mud with lower sedimentation rates. The Paq and C₂₃/C₃₁ ratios increased markedly, which indicates an increased amount of aquatic plant input compared to the last stage. The n-alkane distribution exhibited a two-peak pattern, which supports the Paq and C₂₃/C₃₁ changes. The more negative δD_mean values implied wetter conditions, which corresponded with the sharp increase in the aquatic type of phytolith (Zhang et al., 2019). The wet climate during this period may have favored the growth of aquatic plants. The finer sedimentary conditions of silty mud also indicated the wet climate (Li et al., 2017). The more negative δ¹³C_mean values from the first ten decades implied that the C₃ plants initially grew rapidly, and continuously less negative δ¹³C_mean values demonstrated increasing C₄ plants. The concentration of Cyperaceae pollen peaked during this humid period (Cui et al., 2018), indicating the growth of C₄ plants. The temperature remained high and the warm period continued, where Iw values exhibited a similar trend (Zhang et al., 2019). Herein, it is assumed that this period featured warm and wet conditions. The wet condition did not restrict the growth of C₃ and C₄ plants; however, the long-lasting high temperature decreased C₃ plant growth and contributed to C₄ plant recovery. C₄ plants gradually surpassed C₃ plants before peaking.

Stage III (1440–1650 AD) (dark-blue bar in Figures 5 and 6): The sediment was dark peat with considerably lower sedimentation rates of approximately 0.17 cm/yr. The remarkably increasing δD_mean values indicate a drier condition, and the low P_aq and C₂₃/C₃₁ ratio suggest a decrease in aquatic plant input, which could also be reflected by the low pollen concentration of the wetland herbs (Cui et al., 2018). The one-peak mode of n-alkane distribution reveals that terrestrial higher plants were dominant. The δ¹³C_mean values peaked at approximately 1450 AD and then gradually decreased during this stage. C₃ plants thrived and gradually dominated, whereas C₄ plants suffered. The temperature records showed a significant cooling trend (Liu, 2004; Zhang et al., 2000), which was supported by the decreasing Iw values (Zhang et al., 2019). C₄ plants are significantly influenced by temperature than by precipitation, and therefore, the cooling condition depressed the growth of C₄ plants, although dry condition was suitable for C₄ plants.

Stage IV (1650–1800 AD) (light-blue bar in Figures 5 and 6): The dark peat during this period was deposited at a higher rate. The markedly decreasing δD_mean values reveal the significant transition to a wet condition. The increasing P_aq and C₂₃/C₃₁ ratios indicate the growing amount of aquatic plant input, which was supported by the larger pollen percentage of wetland herbs (Cui et al., 2018). The significantly decreasing δ¹³C_mean values demonstrate a similar trend to that of Stage III, such as the continuous growth of C₃ plants. According to the temperature records, after reaching the lowest anomaly point, the condition remained at a cool level, which was also indicated by the Iw values (Zhang et al., 2019). It is assumed that the low temperature weakened the C₄ plants, whereas the wet conditions favored the thriving of C₃ plants. C₃ plants were dominant again after this period. A wet and cool environment results in coarse particles.

Stage V (1800–2010 AD) (green bar in Figures 5 and 6): The dark peat during this period accumulated at a slightly higher stable rate than Stage IV. The δD_mean values remained negative at a level similar to that of Stage IV. Increasing P_aq and C₂₃/C₃₁ ratios suggested a slight shift to wet conditions, with more aquatic plants in the peat. The lower concentrations of wetland herbs and Iph values exhibit a reverse trend (Cui et al., 2018; Zhang et al., 2019). This contradiction is likely caused by human activities because of the increasing population from 1800 AD. The wetland herbs on the hillsides might have been cleared for rice or tea cultivation, which would have probably decreased original wetland herbs. The δ¹³C_mean values increased slightly when the temperature records implied a shift to warm conditions. The influencing factors of C₃ and C₄ plants cannot be explained because human activities may have begun to influence the evolution of vegetation since approximately 1800 AD.

Combined with temperature records, the curves described herein represent five primary climatic states. These are the warming climate and overall dry condition from 900 to 1125 AD, the warm and wet climate from 1125 to 1450 AD, the cool and dry condition from 1450 to 1650 AD, the cool and wet climate from 1650 to 1800 AD, and the warm and wet climate from 1800 to 2010 AD. Previous studies have shown several climate intervals in the Northern Hemisphere, namely MWP (900–1300 AD), LIA (1350–1850 AD), and PWP (1900 AD to now). In the present study, n-alkane records showed that Stage I and Stage II from 900 to 1450 AD could be a response to MWP. The MWP could be divided into two periods, and the distributions of n-alkanes and δD_mean values revealed a dry-to-wet transformation. Overall, the records exhibited a wetter MWP. Climate changes in Stages III and IV from 1450 to 1800 AD respond to the LIA period, which could also be divided into two sub-climatic conditions, with the former characterized by a drier and cooling condition, and the latter being a considerably wetter and cooling condition. After the LIA, the later part of Stage V beginning in 1800 AD corresponds to the PWP, with an overall warm and wet climate.

Compared with the LIA, the core sections of the MWP and the PWP contain more fine particles (Li et al., 2017), which also supports the warmer and wetter conditions than those of the LIA. The interval between 1450 and 1650 AD had a significantly lower sedimentation rate (Figure 2). Although the n-alkane proxies, indicating a cooler and drier climate during this period would constrain microbial degradation and facilitate the accumulation of organic matter, vegetation types could significantly affect the degradation. The decay rate of vascular peat-forming plants is generally higher than that of Sphagnum species (Schellekens et al., 2015). The C₂₃/C₃₁ ratio and P_aq were low during 1450–1650 AD, indicating the low input of aquatic plants, and the less negative δ¹³C_mean values indicated the increased growth of C₄ sedges. It is assumed that the vegetation type change is the primary cause of the remarkably low sedimentation rate of 1450–1650 AD.

Climatic intercomparison and possible controlling factors

Comparison of paleoclimate records

The changes in climate and vegetation from the current findings, particularly during the universal MWP, LIA, and PWP periods (Figure 5a–c and j), were recorded at other sites or historical records in the East Asia monsoon area. The XYH peatland δD_mean record and C₂₃/C₃₁ ratio indicate centennial-scale fluctuations in the hydroclimate, with an overall wetter MWP, drier LIA, and wetter PWP (Figure 5c). There exist sub-stages in both the MWP and LIA, such as the drier-to-wetter transfer during the MWP and the drier-to-wetter transfer during the LIA. The D–W index (Figure 5g) (Jingyun et al., 2006) calculated from official historical records in the Jian-Nan area (approximately 25°N–31°N) suggested the same trend. However, there are differences, such as the degree of wetter conditions during the early LIA from 1450 to 1650 AD and the level of aridity during the later LIA from 1650 to 1800 AD. This discrepancy is acceptable for the deviation between historical and natural records. This is not the only hydrology reconstruction in the middle–lower Yangtze River area, and although it is a rare high-resolution δD reconstruction from ombrotrophic peatlands covering the MWP and LIA period, we compared the values to proxy archives in the region. The late MWP and late LIA trended toward more depleted speleothem δ¹⁸O from Heshang Cave, and therefore, the early MWP and LIA trended toward more enriched values of δ¹⁸O (Figure 5f) (Hu et al., 2008). Although there are some anomalies because of dating error and resolution effects, the C₂₃/C₃₁ ratio, δD_mean, and δ¹⁸O generally exhibited a similar hydroclimate history. The Poyang Lake δD_C31 of plant wax from lake sediments exhibited a discrepancy, where the lake δD_C31 values (Figure 5h) indicated a wetter condition in the early LIA and a drier condition in the late LIA (Yao et al., 2015). As discussed earlier, δD from plant wax could be affected by the “amount effect,” which would be recorded in the sedimentary n-alkane δD values (Sachse et al., 2012). The study site is nearby Poyang Lake. However, peatland characteristics, such as high relative humidity and relatively singular plant group, deplete the “amount effect” and the effect of vegetation lifeforms, which may lead to the short time difference. In general, plant lipids from sedimentary accumulations tend to average and alleviate these effects (Sachse et al., 2012), and therefore, the robust δD records of peatland and lake cores reveal a wetter MWP and a drier LIA.

Possible controlling factors

Previous studies show that the δ¹⁸O values of stalagmites are synchronized with the intensity of the East Asian Summer Monsoon (EASM), which controls the amount of precipitation (An, 2000; Ding et al., 2008; Zhang et al., 2011; Zheng et al., 2017). The humidity variations illustrated by n-alkane proxies and δD correspond to δ¹⁸O, and therefore, δD could be an indicator of EASM intensity, which is also significantly influenced by the El Niño Southern Oscillation (ENSO) (An, 2000; Chen et al., 2010, 2015; Oppo et al., 2009;). The West Pacific Subtropical High moves toward the southwest causing abundant precipitation during El Niño years and moves toward the northeast during La Niña years. Cold sea surface Temperatures (SSTs) in the eastern equatorial Pacific (La Niña) are correlated with excessive precipitation in Southeast Asia, whereas a warm eastern tropical Pacific (El Niño) causes drought conditions (Chen et al., 2015; Xu et al., 2016). In this study, the East Equatorial Pacific (EEP) SST (Figure 6g) was low during the early MWP, then swiftly turned high, and later returned to low values. The n-alkane data identified progressively wetter El Niño conditions from the MWP and drier La Niña conditions at the end of the MWP. The cool and dry EEP conditions associated with La Niña since 1450 AD are shown by n-alkane data, and El Niño conditions since 1800 AD later caused a slightly wetter climate. The SST and n-alkanes data showed that El Niño gradually predominated since 1800 AD. The generally warm-humid MWP period corresponds to El Niño-like conditions (relatively higher yet cooling SST of the EEP), whereas the cool-drier LIA period corresponds to La Niña-like conditions (relatively colder SST of the EEP). The Niño 3.4 index emphasizes the El Niño-like conditions in the MWP and the PWP and conspicuous La Niña-like conditions in the LIA.

During the warm period, the C₃ plant-dominated scenario had fewer graminoids and woody angiosperms, which typically have lower CPIs. Furthermore, the more productive bacterial diagenesis under warmer conditions could also lead to the lower CPIs (Naafs et al., 2019). These two factors lead to low CPIs. In contrast, during the cool period, the abundance of C₄ plants brings more graminoids and woody angiosperm plants, and the less productive diagenesis leads to higher CPIs. The CPI curve indicates that the centennial warm MWP and PWP experienced partially stronger solar irradiance, whereas the cooler LIA experienced weaker irradiance.

The Mount Xishan study area was not affected by human activities until the Song Dynasty (960–1279 AD), following which time war caused increased migration (Xu, 2000). North–south differences in the environment (northern drought and humid south) encouraged invaders from Mongolia to move toward southern China and compete for resources. Migration contributed to the population of Jiang Xi Province (Nan Chang County Annals) and lasted until the Yuan Dynasty (1271–1368 AD) when the southern China experienced a warm period (Yan et al., 2015). The warm and humid climate in the middle Yangtze River catchment encouraged population development. Although the stable society and agriculture-oriented policies supported by the Yuan Dynasty were suitable for development, the population decreased substantially, likely because of the cool and dry climate during the LIA. This trend may have led to the fall of the Ming Dynasty (1368–1644 AD). After the LIA, the warmer and more humid climate enabled a significant increase in population after 1750 AD.

Conclusions

In this study, a biomarker approach was applied to investigate the past 1200 years of climate variability from a peat core in the XYH peatland in southeast China. The input changes of the aquatic plants from n-alkane distribution patterns and the parameters (P_aq and C₂₃/C₃₁ ratios) were reconstructed. Furthermore, the paleohydrological changes based on the δD from plant wax with the n-alkane changes and inferred vegetation changes from δ¹³C were qualitatively reconstructed. During the MWP (900–1450 AD), P_aq, C₂₃/C₃₁ ratios, and δD generally inferred that the hydroclimate in the middle Yangtze River is wet and C₄ vegetation increased under the dominance of C₃ vegetation. In contrast, during the LIA (1450–1800 AD), a shift to relatively drier conditions occurred, and C₃ vegetation generally increased. In the PWP (1800–2010 AD), the condition again became wetter compared to the LIA. This study also revealed that there are two intervals between the MWP and the LIA. The results are generally consistent with previous paleoclimatic reconstructions nearby and in situ but exhibit different hydroclimatic patterns to those from the southwest and northeast China in the monsoon areas (Chen et al., 2015).

In this study, it is proposed that the changes in ENSO and the EASM may have played key roles in producing hydroclimatic contrasts during the MWP and the LIA. The drier LIA was likely caused by the weaker EASM and La Niña conditions, whereas the wetter MWP and PWP resulted from the opposite conditions. Climatic conditions play an important role in social development. It is suggested that the up-and-down trend of the population during the MWP and the LIA was affected by the different water-heat conditions. With the development of instruments and fertilizers, the impact of human activities might have affected the local vegetation and microclimate from 1800 AD. However, further research is required to confirm this hypothesis

Research Data

sj-xls-1-hol-10.1177_09596836211025966 – Research Data for Biogeochemical evidence for environmental and vegetation changes in peatlands from the middle Yangtze river catchment during the medieval warm period and little ice Age

Research Data, sj-xls-1-hol-10.1177_09596836211025966 for Biogeochemical evidence for environmental and vegetation changes in peatlands from the middle Yangtze river catchment during the medieval warm period and little ice Age by Jia Sun, Chunmei Ma, Bin Zhou, Jiawei Jiang and Cheng Zhao in The Holocene

This article is distributed under the terms of the Creative Commons Attribution 4.0 License (https://creativecommons.org/licenses/by/4.0/) which permits any use, reproduction and distribution of the work without further permission provided the original work is attributed as specified on the SAGE and Open Access pages (https://us.sagepub.com/en-us/nam/open-access-at-sage).

Footnotes

Acknowledgements

We thank Chaohao Ling and Guangjiu Ling for their assistance with field sampling work and the Nanjing Institute of Geography & Limnology, Chinese Academy of Sciences for the isotope test.

Availability of data and material

Datasets archiving for this research is underway. We have uploaded our data on Pangaea.de, yet they are being checked and processed. A copy of this data could be found in the “Supporting Information” file.

Funding

The author(s) disclosed receipt of the following financial support for the research, authorship, and/or publication of this article: This work was supported by the National Natural Science Foundation of China (grant numbers 41371202, 41671196, and 41977378), Major Social Science Project (20&ZD247), and the Global Change Program of the National Key Research and Development Program of China (grant number 2016 YFA0600501).

ORCID iD

Chunmei Ma

References

(2000) The history and variability of the East Asian paleomonsoon climate. Quaternary Science Reviews 19(1): 171–187.

Baas

Pancost

van Geel

, et al. (2000) A comparative study of lipids in Sphagnum species. Organic Geochemistry 31(6): 535–541.

Bard

Raisbeck

Yiou

, et al. (2000) Solar irradiance during the last 1200 years based on cosmogenic nuclides. Tellus B 52(3): 985–992.

Barnes

(1978) Organic compounds in lake sediments. In: Lerman

(ed.) Lakes: Chemistry, Geology, Physics. New York: Springer New York, pp.127–152.

Blaauw

Christeny

(2011) Flexible paleoclimate age-depth models using an autoregressive gamma process. Bayesian Analysis 6(3): 457–474.

Bray

Evans

(1961) Distribution of n-paraffins as a clue to recognition of source beds. Geochimica et Cosmochimica Acta 22(1): 2–15.

Bureau

(2015) China Wetland Resources (Jiangxi Volume). Beijing: China Forestry Publishing House.

Castañeda

Mulitza

Schefuß

, et al. (2009) Stable oxygen and carbon isotopes ratios of sediment core GeoB9528-3. Supplement to: Castañeda, IS et al. (2009): Wet phases in the Sahara/Sahel region and human migration patterns in North Africa. Proceedings of the National Academy of Sciences of the United States of America 106(48): 20159–20163.

Chen

Feng

, et al. (2015) Hydroclimatic changes in China and surroundings during the medieval climate anomaly and little ice age: Spatial patterns and possible mechanisms. Quaternary Science Reviews 107: 98–111.

10.

Chen

F-H

Chen

J-H

Holmes

, et al. (2010) Moisture changes over the last millennium in arid central Asia: A review, synthesis and comparison with monsoon region. Quaternary Science Reviews 29(7): 1055–1068.

11.

China ECfVMo (2001) Vegetation Atlas of China. Beijing: Beijing: Science Press.

12.

Collister

Rieley

Stern

, et al. (1994) Compound-specific δ 13C analyses of leaf lipids from plants with differing carbon dioxide metabolisms. Organic Geochemistry 21(6): 619–627.

13.

Conroy

Overpeck

Cole

(2010) El Niño/Southern Oscillation and changes in the zonal gradient of tropical Pacific sea surface temperature over the last 1.2 ka. PAGES News 18: 32–34.

14.

Cranwell

(1984) Lipid geochemistry of sediments from Upton Broad, a small productive lake. Organic Geochemistry 7(1): 25–37.

15.

Cranwell

Eglinton

Robinson

(1987) Lipids of aquatic organisms as potential contributors to lacustrine sediments—II. Organic Geochemistry 11(6): 513–527.

16.

Cui

Zhao

, et al. (2018) Pollen records of the little ice age humidity flip in the middle Yangtze River catchment. Quaternary Science Reviews 193: 43–53.

17.

D’Arrigo

Wilson

Jacoby

(2006) On the long-term context for late twentieth century warming. Journal of Geophysical Research: Atmospheres 111(D3): D006352.

18.

Ding

Wang

Sun

(2008) Inter-decadal variation of the summer precipitation in East China and its association with decreasing Asian summer monsoon. Part I: Observed evidences. International Journal of Climatology 28(9): 1139–1161.

19.

Eglinton

Hamilton

(1967) Leaf epicuticular waxes. Science 156(3780): 1322–1335.

20.

Emile-Geay

Cobb

Mann

, et al. (2013) Estimating central equatorial pacific SST variability over the past millennium. Part II: Reconstructions and implications. Journal of Climate 26: 2329–2352.

21.

Feakins

Ponton

, et al. (2019) Biomarkers reveal abrupt switches in hydroclimate during the last glacial in southern California. Earth and Planetary Science Letters 515: 164–172.

22.

Ficken

Swain

, et al. (2000) An n-alkane proxy for the sedimentary input of submerged/floating freshwater aquatic macrophytes. Organic Geochemistry 31(7): 745–749.

23.

Gałka

Tanţău

Feurdean

(2017) Plant succession in a peatland in the Eastern Carpathian Mts. (CE Europe) during the last 10,200years: Implications for peatland development and palaeoclimatic research. Review of Palaeobotany and Palynology 244: 203–216.

24.

Holzhauser

Magny

Zumbuühl

(2005) Glacier and lake-level variations in west-central Europe over the last 3500 years. The Holocene 15(6): 789–801.

25.

Hou

D’Andrea

Huang

(2008) Can sedimentary leaf waxes record D/H ratios of continental precipitation? Field, model, and experimental assessments. Geochimica et Cosmochimica Acta 72(14): 3503–3517.

26.

Henderson

Huang

, et al. (2008) Quantification of Holocene Asian monsoon rainfall from spatially separated cave records. Earth and Planetary Science Letters 266(3): 221–232.

27.

Huang

Pancost

Jiantao

, et al. (2018) Response of carbon cycle to drier conditions in the mid-Holocene in central China. Nature Communication 9: 1369.

28.

Jingyun

Wang

W-C

, et al. (2006) Precipitation variability and extreme events in eastern China during the past 1500 years. Terrestrial, Atmospheric and Oceanic Sciences 17: 579.

29.

Kayastha

Harrison

(2008) Changes of the equilibrium-line altitude since the Little Ice Age in the Nepalese Himalaya. Annals of Glaciology 48: 93–99.

30.

Lichtfouse

Elbisser

Balesdent

, et al. (1994) Isotope and molecular evidence for direct input of maize leaf wax n-alkanes into crop soils. Organic Geochemistry 22(2): 349–351.

31.

, et al. (2017) A preliminary study of the climate change since 2ka archived by a peat from Yuhua Moutain in the middle Jiangxi provience. Quaternary Research 37(3): 548–12 (in Chinese).

32.

Liu

(2004) Climate Process and Change. Beijing: Science Press of China, pp. 236–238.

33.

Liu

(2020) Leaf wax n-alkane carbon isotope values vary among major terrestrial plant groups: Different responses to precipitation amount and temperature, and implications for paleoenvironmental reconstruction. Earth-Science Reviews 202: 103081.

34.

Lorrey

Fauchereau

Stanton

, et al. (2014) The Little Ice Age climate of New Zealand reconstructed from Southern Alps cirque glaciers: A synoptic type approach. Climate Dynamics 42(11): 3039–3060.

35.

Mann

Bradley

Hughes

(1999) Northern hemisphere temperatures during the past millennium: Inferences, uncertainties, and limitations. Geophysical Research Letters 26(6): 759–762.

36.

Matthews

Briffa

(2016) The ‘little ice age’: Re-evaluation of an evolving concept. Geografiska Annaler: Series A, Physical Geography 87(1): 17–36.

37.

Naafs

BDA

Inglis

Blewett

, et al. (2019) The potential of biomarker proxies to trace climate, vegetation, and biogeochemical processes in peat: A review. Global and Planetary Change 179(AUG.): 57–79.

38.

Nichols

Booth

Jackson

, et al. (2006) Paleohydrologic reconstruction based on n-alkane distributions in ombrotrophic peat. Organic Geochemistry 37(11): 1505–1513.

39.

Nichols

Walcott

Bradley

, et al. (2009) Quantitative assessment of precipitation seasonality and summer surface wetness using ombrotrophic sediments from an Arctic Norwegian peatland. Quaternary Research 72(3): 443–451.

40.

Nott

Xie

Avsejs

, et al. (2000) n-Alkane distributions in ombrotrophic mires as indicators of vegetation change related to climatic variation. Organic Geochemistry 31(2–3): 231–235.

41.

Normile

(1997) Yangtze seen as earliest rice site. Science 275(5298): 309.

42.

Oppo

Rosenthal

Linsley

(2009) 2,000-year-long temperature and hydrology reconstructions from the Indo-Pacific warm pool. Nature 460(7259): 1113–1116.

43.

Ortiz

Gallego

JLR

Torres

, et al. (2010) Palaeoenvironmental reconstruction of Northern Spain during the last 8000 cal yr BP based on the biomarker content of the Roñanzas peat bog (Asturias). Organic Geochemistry 41(5): 454–466.

44.

Pang

Zhou

, et al. (under review) Alkane variation in peat reveals palaeohydrological changes since the Little Ice Age in southeastern China. Palaeogeography, Palaeoclimatology, Palaeoecology.

45.

Poynter

Farrimond

Robinson

, et al. (1989) Aeolian-derived higher plant lipids in the marine sedimentary record: Links with palaeoclimate. In: Leinen

Sarnthein

(eds.) Paleoclimatology and Paleometeorology: Modern and Past Patterns of Global Atmospheric Transport. Dordrecht: Springer Netherlands, pp.435–462.

46.

Routh

Hugelius

Kuhry

, et al. (2014) Multi-proxy study of soil organic matter dynamics in permafrost peat deposits reveal vulnerability to climate change in the European Russian Arctic. Chemical Geology 368: 104–117.

47.

Sachse

Billault

Bowen

, et al. (2012) Molecular paleohydrology: Interpreting the hydrogen-isotopic composition of lipid biomarkers from photosynthesizing organisms. Annual Review of Earth and Planetary Sciences 40(1): 221–249.

48.

Sachse

Radke

Gleixner

(2006) δD values of individual n-alkanes from terrestrial plants along a climatic gradient – Implications for the sedimentary biomarker record. Organic Geochemistry 37(4): 469–483.

49.

Schellekens

Bindler

Martínez-Cortizas

, et al. (2015) Preferential degradation of polyphenols from Sphagnum – 4-Isopropenylphenol as a proxy for past hydrological conditions in Sphagnum-dominated peat. Geochimica et Cosmochimica Acta 150: 74–89.

50.

Schittek

Kock

Lücke

, et al. (2016) A high-altitude peatland record of environmental changes in the NW Argentine Andes (24 ° S) over the last 2100 years. Climate of the Past 12(5): 1165–1180.

51.

Wang

Zheng

Man

, et al. (2017) Branched GDGT-based paleotemperature reconstruction of the last 30,000years in humid monsoon region of Southeast China. Chemical Geology 463: 94–102.

52.

(2000) Interaction between the economic development and ecological environment in Jiangxi history. Agricultural Archaeology 3: 110–124. (in Chinese)

53.

Jia

, et al. (2016) Geographic distribution of archaeological sites and their response to climate and environmental change between 10.0-2.8 ka BP in the Poyang Lake Basin, China. Journal of Geographical Sciences 26(5): 603–618.

54.

Yan

Zhang

Wang

, et al. (2015) Simulated warm periods of climate over China during the last two millennia: The Sui-Tang warm period versus the Song-Yuan warm period. Journal of Geophysical Research: Atmospheres 120(6): 2229–2241.

55.

Yao

Yang

Liu

, et al. (2015) Hydrological changes of the past 1400years recorded in D of sedimentary n-alkanes from Poyang Lake, southeastern China. Holocene 25(7): 94–99.

56.

Zhang

Chen

Holmes

, et al. (2011) Holocene monsoon climate documented by oxygen and carbon isotopes from lake sediments and peat bogs in China: A review and synthesis. Quaternary Science Reviews 30(15): 1973–1987.

57.

Zhang

Fang

Ren

(2000) Global Change. Beijing: Higher Education Press, 431.

58.

Zhang

, et al. (2019) Climatic controls on peat swamp formation and evolution since 1300 year BP as recorded by phytoliths in the Xishan Mountains, Jiangxi Province, China. Palaeogeography, Palaeoclimatology, Palaeoecology 522: 76–88.

59.

Zhao

Piperno

(2000) Late Pleistocene/Holocene environments in the middle Yangtze river valley, China and rice (Oryza sativa L.) domestication: The phytolith evidence. Geoarchaeology 15(2): 203–222.

60.

Zheng

Wang

, et al. (2006) Precipitation variability and extreme events in eastern China during the past 1500 years. Terrestrial Atmospheric and Oceanic Sciences 17(3): 579–592.

61.

Zheng

, et al. (2017) Observed, reconstructed, and simulated decadal variability of summer precipitation over eastern China. Journal of Meteorological Research 34(1): 49–60.

62.

Zheng

Zhou

Xie

, et al. (2009) Comparative study of n-alkane biomarker and pollen records: An example from southern China. Science Bulletin 54(6):1065–1072.

63.

Zhou

Xie

Meyers

, et al. (2005) Reconstruction of late glacial and Holocene climate evolution in southern China from geolipids and pollen in the Dingnan peat sequence. Organic Geochemistry 36(9): 1272–1284.

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