2017, Vol.37
② 山东师范大学地理与环境学院, 济南 250014;
③ 中国科学院大学, 北京 100049)
根据植物光合作用的途径差异，陆生高等植物主要分属于Calvin-Benson(C3)植物和Hatch-Slack(C4)植物两类。现代C3植物包含所有的木本植物、大多数灌木和一些耐寒的草本植物，其稳定碳同位素(δ13C)组成一般在-34 ‰ ~-22 ‰之间，平均约为-27 ‰；绝大部分C4植物是草本植物，主要由藜科、禾本科和莎草科的部分植物组成，其δ13C分布在-16 ‰ ~-10 ‰之间，平均约为-13 ‰[1, 2]。C3和C4植物之间不仅具有显著的碳同位素差异，而且在不同的环境条件下具有不同的生长优势。现代植物分布研究表明，通常情况下低温、湿润和较高的大气二氧化碳浓度(pCO2)的环境有利于C3植物的生长，而高温、干燥和较低的pCO2的环境更适合以C4植物占优势的群落的形成[3]。因此，黄土/古土壤、湖泊和海洋等地质载体中有机质的δ13C组成被广泛应用于地质历史时期古植被与古环境的重建[4~13]。
在末次冰消期，大气温室气体浓度和全球温度显著升高，大陆冰盖消退，是研究全球变化及其环境效应的重要阶段[14]。最初的湖泊沉积研究发现，末次冰消期在东非肯尼亚地区C4植物逐渐减少，在全新世形成了以C3植物为主导的群落，这可能是由于大气pCO2逐渐升高造成的[15]。但是随后的研究却表明末次冰盛期以来湖泊沉积物中正构烷烃单体δ13C的变化存在显著的区域性，其相应的C3/C4植物组成变化主要受气候条件的控制[16]。在我国，末次冰盛期以来C3/C4植物组成变化的研究主要集中在黄土高原地区，结果显示该区域末次冰期植被类型以C3植物为主导，C4植物的相对丰度在全新世迅速增加，也表明pCO2不是控制黄土高原地区植被组成的直接因素[8, 17, 18]。在我国低纬地区，末次冰盛期以来C3/C4植物组成变化的记录还比较缺乏，大部分来自于湖泊和海洋的沉积记录表明末次冰盛期我国南方地区分布着比较丰富的C4植物[7, 19, 20]，而云南泸沽湖的长链正构烷烃单体δ13C记录却显示区域C4植物的相对丰度在末次冰盛期却低于早全新世[21]。这些研究虽然在C3/C4植物相对丰度变化及其驱动机制上取得了显著的进展，但是这些记录的时间分辨率较低，注重于冰期-间冰期时间尺度上的植被转换过程，对于具体的研究区域，特别是低纬地区，C3/C4植物相对丰度的变化如何响应于气候突变事件和温室气体浓度变化，仍然是一个亟待解决的问题。因此，阐明区域植被变化的主要驱动因素需要更多的具有可靠年代、分辨率较高的古植被和古环境记录，这对人类适应未来气候变化和温室气体浓度持续上升可能导致的陆地生态系统变化具有重要意义。
在地质构造作用下，我国西南地区分布着较多的断陷深水湖，其形成年代在上新世末至第四纪初，其沉积物具有沉积较为连续、速率较高的特点[22]。然而，湖泊沉积物的有机碳通常由陆生植物碎屑和湖泊水生生物共同组成，此外，沉水植物的δ13C分布在-50 ‰ ~-11 ‰之间，难以利用全有机质的δ13C来重建陆生高等植被的演化过程[23, 24]。黑碳或元素碳，是由生物质和化石燃料不完全燃烧产生的一种含碳物质的连续统一体，包括木炭/焦炭、烟炱和石墨态物质等[25, 26]。元素碳在经历了一系列地质营力的搬运后最终会广泛分布于土壤、河流、湖泊和海洋等沉积载体中。由于元素碳的惰性特征，其沉积后受光化学反应和微生物作用的影响很小，可以长期存在于地质环境中。在燃烧过程中，由于生物质的化学组成不同，产生的同位素分馏效应也有所差异，但是元素碳δ13C(δ13CEC)相对于植物体δ13C的变化总体较小[6, 25]。目前，越来越多的研究开始利用δ13CEC重建区域的植被演替与气候变化[6, 7, 27~30]。本研究拟通过对云南洱海柱状沉积物中δ13CEC的分析，重建该区域末次冰盛期以来的植被变化历史，为进一步探讨C3/C4植被在千年尺度上对气候与温室气体浓度变化的响应提供新的依据。
2 研究区概况洱海(25°25′~26°16′N，99°32′~100°27′E)属高原断陷湖泊，形成于更新世早期，地处云贵高原的西北部，横断山脉的东部边缘[31]。湖泊水面海拔1973m，湖盆呈南北向展布，面积约150km2，流域面积2785km2。湖泊最大水深20.7m，平均水深10.2m。湖水主要依靠入湖径流和降水补给，主要入湖河流有弥苴河/罗时江和波罗江，出流河仅西南端的西洱河(图 1)，最终汇入澜沧江。流域基岩以沉积岩和变质岩为主，地带性土壤为红壤。湖泊西部为点苍山，海拔最高为4100m左右，东部为丘陵地带。区域自然植被的垂直分带较为明显，但由于人类活动的干扰，山麓坡地植被主要由云南松(Pinus yunnanensis)和落叶栎(Deciduous Quercus)组成，冲积扇和湖滨地带均改造成农田等人为景观。流域气候属低纬高原亚热带季风气候，四季干湿分明，据大理气象站多年记录，该区年均气温15℃，年均降水量1060mm，降水主要集中在5~10月，占全年降水的85 %以上[31]。
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图 1 云南洱海地理位置及采样点 Fig. 1 Location of Lake Erhai, Yunnan and the coring site |
2012年5月，在水深13.5m处(25°51.02′N，100°11.02′E)用采样平台钻取了一根长684cm的柱状岩芯，运回实验室后以1cm间距分样，在4℃下低温保存，直至样品分析。定年材料为柱状沉积物中不同深度留存的11个植物残体和2个炭屑样品，植物残体测试在美国Beta AMS14C实验室进行，而炭屑样品在波兰波茨南大学AMS14C实验室进行。所有的AMS14C年龄都基于IntCal13数据库的Calib 7.10校正程序进行了校正[32]，并根据Bacon程序中贝叶斯模型[33, 34]建立沉积岩芯的深度-年龄模型。
本研究以4cm的间距，对171个沉积物样品进行了δ13CEC分析。元素碳的提取采用湿化学氧化法[35]。具体的步骤如下：首先称取1~2g研磨过的沉积物干样，加入3mol/L的盐酸(HCl)，以去除沉积物中可能存在的碳酸盐；加入浓度比为10：1的HF与HCl的混合液，以去除硅酸盐晶格中的碳酸盐及部分硅酸盐矿物；再加入10mol/L的HCl去除上步反应可能产生的CaF2沉淀；加入0.1mol/L的KOH溶液，去除沉积物中的腐殖酸；将剩余样品烘干、研磨粉碎，取0.1~0.2g样品放入玻璃试管中，加入0.2mol/L的K2Cr2O7的硫酸溶液(硫酸浓度为2mol/L)，在55℃水浴锅中反应60小时，以去除易降解的有机质和干酪根，在60℃烘箱中烘干。处理后样品中所含的有机碳即认为是难降解的元素碳。δ13CEC的测试仪器为Thermo Delta Plus同位素质谱仪，测试在中国科学院南京地理与湖泊研究所湖泊与环境国家重点实验室完成。δ13C用VPDB标准方式表示，分析的精度优于0.2 ‰。
4 结果通过贝叶斯模型获得的年代-深度关系如图 2所示，岩芯底部加权平均年龄为19.4cal.ka B.P.，沉积速率变化较大，在0.01~0.37cm/年之间。每1cm年代的95 %置信范围的平均值为1092年，底部最大达到5170年。
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图 2 洱海沉积岩芯基于Bacon程序[33]中贝叶斯模型建立的年代-深度关系灰色虚线代表 95 %的置信区间 Fig. 2 Age-depth model for the Lake Erhai sediment core produced by Bacon software[33]. Gray dotted lines indicate the 95 % confidence range |
洱海沉积物δ13CEC值随时间变化特征如图 3a所示，其变化范围在-30.0 ‰ ~-20.8 ‰之间，平均约为-25.9 ‰，全新世δ13CEC值整体偏负于晚冰期。其主要变化特征可以分为4个阶段：在14.7cal.ka B.P.之前，δ13CEC值较高且相对稳定，在-22.6 ‰ ~-20.8 ‰之间，平均约为-21.8 ‰；在14.7~11.5cal.ka B.P.期间，δ13CEC值逐渐偏负，平均约为-25.6 ‰；在早中全新世(11.5~5.8cal.ka B.P.)，δ13CEC值较低，平均约为-27.4 ‰；此后δ13CEC值略有升高，在-23.3 ‰ ~-28.3 ‰之间，平均约为-26.2 ‰。此外，δ13CEC值在16.1~15.3 cal.ka B.P.、12.3~11.5cal.ka B.P.、8.2~7.1cal.ka B.P.和5.7~4.7cal.ka B.P.显著偏正。
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图 3
末次冰盛期以来洱海沉积物δ13CEC与其他古环境记录的对比
(a)洱海沉积物δ13CEC；(b)董哥洞石笋δ18O记录[58]；(c)腾冲青海 中值粒径[63]；(d)南极EPICA Dome C冰芯大气pCO2[41, 42] Fig. 3 Comparson of the δ13CEC record of Lake Erhai with other paleoenvironmental proxies since the Last Glacial Maximum. (a)δ13CEC record of Lake Erhai; (b)δ18O record from Dongge Cave[58]; (c)Median grain-size record from Lake Tengchongqinghai[63]; (d)CO2 record from the EPICA Dome C ice core, Antarctica[41, 42] |
一般来说，燃烧过程中的同位素分馏效应、大气CO2浓度及其碳同位素、温度、降水和陆地植被组成是可能影响地质载体中δ13CEC的主要因素。在燃烧过程中，优先分解热稳定性较差的半纤维素、糖类、氨基酸和果胶等化合物，不完全燃烧的产物中木质素、纤维素和脂类等化合物的比例相应增加，δ13CEC与原植物体碳同位素组成存在一定的差异，且多为负值[25, 36, 37]。燃烧实验表明，C3植物燃烧前后δ13C的变化不大，在-3 ‰ ~3 ‰之间，平均为-0.3 ‰；而C4植物燃烧生成的元素碳受物种、燃烧温度和时间的影响，其δ13C较原植物体的变化最大可以偏负10 ‰[6]。一方面可能是由于C4植物相对于C3植物含有更高比例的半纤维素等热稳定性较差的化合物，另一方面可能是C4植物植硅体中封存着δ13C要明显偏负于植物的有机质，而C3植物植硅体封存的有机质δ13C与植物非常接近[25, 36, 37]。C4植物不完全燃烧产生的元素碳同位素分馏效应更加显著，但平均仅为-1.7 ‰，因此，在考虑燃烧过程产生的同位素分馏效应的前提下，δ13CEC指标可以用来反映燃烧前植物的δ13C组成情况[6]。
理论上讲，pCO2的升高，可以导致C3植物的δ13C偏负，这是因为在CO2浓度较低时，植物叶肉细胞间CO2与周围大气环境的CO2气压差减低，C3植物的光合作用受到限制，而植物为了减缓CO2浓度降低的胁迫，气孔收缩，进入叶肉细胞的CO2减少，减弱了羧化过程中同位素分馏作用，而C4植物比C3植物具有更高的CO2利用效率，对C4植物δ13C组成的影响较弱[1, 38, 38]。另外，由于大量的化石燃料的使用，自工业革命以来，大气CO2的δ13C比工业革命前的偏负约1.1 ‰，也会直接影响到植物的δ13C组成[1, 40]。在末次盛冰期-全新世的转变过程中，大气pCO2大幅升高，从185ppm上升到265ppm(图 3d)，而自工业革命以来，pCO2上升了70ppmv[41, 42]，将使C3植物的δ13C分别偏负约2 ‰和1.2 ‰[39]。因此，洱海沉积岩芯从末次冰盛期至工业革命前C3植物的δ13C与现代植被的同位素组成具有显著的差异：末次冰盛期C3植物δ13C的平均值约偏正4.3 ‰，而全新世C3植物δ13C的平均值约偏正2.3 ‰。
此外，温度和降水的变化对C3植物的δ13C组成具有显著的影响[43]。在我国北方，Liu等[44]研究发现，我国西北干旱-半干旱区C3植物δ13C组成与降水量负相关，降水量每增加100mm，植物δ13C偏负约1.1 ‰；而Ma等[45]对该区的研究结果显示降水量每增加100mm，植物δ13C偏负约0.6 ‰；Wang等[46]在黄土高原的研究发现C3植物δ13C随降水量的变化仅为-0.49 ‰ /100mm。以上结果均表明当降水量减少，土壤含水量和有效湿度降低，C3植物为了减少水分的蒸发，会关闭一些气孔，导致气孔通导系数减小，从而引起植物叶内CO2浓度降低，使光合作用产物的δ13C偏重[43]。除了降水量外，温度等气候因素对土壤含水量和有效湿度也有较显著的影响，通常情况下C3植物的δ13C与温度正相关，即温度升高，δ13C偏重[47, 48]。Wang等[49]研究了我国北方400mm等降水线C3植物δ13C与温度的关系，结果表明，植物的δ13C随着温度升高而偏重，气温每升高1℃，δ13C偏重0.104 ‰；藜(Chenopodium album)、平车前(Plantago depressa)和桃叶蓼(Polygonum persicaria)3种C3植物的δ13C也均显示出与温度很好的正相关性。表明温度的降低，减弱了植物的蒸腾作用，并减少了土壤水分的蒸发，促进了C3植物气孔的打开，而使其δ13C变轻[49]。由此可见，C3植物δ13C与环境因子之间具有显著的定性关系，不同的物种对环境因子的敏感程度也不相同。虽然研究区目前仍然缺乏可靠的高分辨率的古气候定量重建资料，但是相关研究表明末次冰盛期以来青藏高原东南缘的温度变化在2~3℃，降水变化幅度也较小[50, 51]。气候变化引起的C3植物光合作用过程中同位素分馏效应可能不是影响洱海沉积物δ13CEC大幅变化的主要因素。
最有可能的解释为洱海沉积物δ13CEC的变化主要反映了区域C3/C4植物相对丰度的变化。洱海沉积物δ13CEC结果表明，在14.7cal.ka B.P.之前，洱海地区分布着比较丰富的C4植物，并在15.7cal.ka B.P.前后达到最大丰度；在14.7~12.3cal.ka B.P.期间，洱海地区的C4植物含量逐渐减少，C3植物含量持续增加；在12.3~11.5cal.ka B.P.期间，C4植物相对丰度明显增加；在全新世，晚全新世C4植物的相对丰度要显著高于早中全新世，并且存在8.2~7.1cal.ka B.P.和5.7~4.7cal.ka B.P.两个持续近千年的扩张阶段。在长时间尺度上，这与我国南方地区已有的C3/C4植物演替记录基本一致(图 4)[6, 19, 20]。云南腾冲青海的δ13CEC研究显示，末次冰盛期是C4植物的近2万年来最繁盛时期，在早中全新世C3植物达到鼎盛[7]；我国南海北部MD05-2905孔沉积物长链正构烷烃单体δ13C从末次冰盛期以来逐渐偏负，表明C4植物在冰期向全新世转变过程中逐渐减少[19]；雷州半岛的泥炭沉积研究表明该区域在MIS2阶段C4植物比较丰富，进入全新世后C3植物迅速扩张[20]；孟加拉湾SO188-342KL孔沉积物长链正构烷烃单体δ13C也反映了C4植物在末次冰盛期最为丰富[52]。另外，Rao等[53]整理了多个低纬地区的湖泊和海洋沉积有机质或长链正构烷烃单体δ13C记录，也发现末次冰期至全新世，C4植物的相对丰度下降，例如：非洲肯尼亚山脉的Sacred湖、乞力马扎罗山脉的Challa湖和赤道非洲的Malawi湖。然而这些记录局限于较低的分辨率，未能显示出C3/C4植物的千年尺度波动规律。
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图 4 低纬地区末次冰盛期以来C3/C4植物演化的对比 (a)洱海沉积物δ13CEC；(b)腾冲青海沉积物δ13CEC[7]；(c)孟加拉湾SO188-342KL钻孔沉积物长链正构烷烃单体δ13C[52]；(d)我国南海北部MD05-2905钻孔沉积物长链正构烷烃单体δ13C[19] Fig. 4 Comparison of the vegetation proxy for Lake Erhai (a) since the Last Glacial Maximum with other records from the low latitudes: (b)δ13CEC values from Lake Tengchongqinghai in Southwest China[7]; (c)δ13C of specific n-alkanes data from Core SO188-342KL in the Bay of Bengal[52]; (d)δ13C of specific n-alkanes data from Core MD05-2905 in the northern South China Sea[19] |
在全球范围内，60°N以北地区仅发现少数几种C4植物，46°S以南地区则没有发现任何C4植物种属，而且这两个区域的C4植物的相对丰度接近于零，而干湿季节明显且具有较高温度的环境更适合C4植物的生长，例如低纬地区的稀树草原，基本由C4植物组成[3]。在我国，目前已经对C3和C4植物的分布及其δ13C特征与气候因子的关系进行了较为系统的研究[54~57]。Rao等[54]通过表土有机质δ13C和正构烷烃单体δ13C研究发现在我国东部，31°~40°N之间的中纬度地区的水热组合是最适合C4植物生长的环境，其中年降水量在500~1200mm，年均温在12℃以上时，以C4植物占优势；该区域以北很可能由于温度过低，超过了C4植物生长的阈值而不适合C4植物的生长，而在该区域以南，过多的降水更适合木本植物的生长，密闭的林地环境限制了C4植物的扩张。在我国的黄土高原地区，白羊草(Bothriochloa ischaemum)、狗尾草(Setaria viridis)和马唐(Digitaria sanguinalis)是最主要的C4植物，温暖湿润的夏季是其主要生长期，丰度从东南向西北方向逐渐降低，而灌木的比例逐渐增加[55, 56]。在青藏高原地区，Wang等[57]共采集了158种植物标本，却仅发现8种C4植物，而且生物量极低，特殊的低温环境明显限制了C4植物的生长。因此生长季的温度和降水量都对现代C4植物的生长具有重要的影响。
从理论上说，末次冰盛期气温降低可能限制C4植物的生长，例如在我国的黄土高原地区，冰期C4植物显著退缩而C3植物相应地扩张[8, 17, 18]。但是末次冰盛期C4植物在洱海等我国的南方地区表现出扩张趋势，与黄土高原地区C3/C4植物相对丰度的变化相反。研究表明，我国西南地区的降温幅度可能在2~3℃[50, 51]，即云南洱海地区当时的年均温在12~13℃，基本可以满足C4植物的生长需要，表明温度的变化对低纬地区C3/C4植物相对丰度的影响强度要显著低于中高纬度。而夏季风降水的减少可能是C4植物显著扩张的主要原因。西南地区的石笋氧同位素(δ18O)在末次冰盛期显著偏正，表明夏季风减弱，气候比较干旱[58]。虽然石笋δ18O受到水汽来源、降水的季节分布、雨量效应和输送距离等多种环境因子的影响，但是腾冲青海、天才湖和泸沽湖的孢粉记录均显示末次冰盛期西南地区的草本植物较为丰富，森林覆盖度较低，降水较少[59~62]。在开阔的环境下，较好的光照条件可以促进C4植物的生长，成为草本植物重要组成部分。另外，末次冰盛期期间，冰芯记录表明pCO2要比现代低140 ppmv[42]，在此环境下，C4植物相比C3植物具有更高的光合作用效率，也更有利于C4植物的生长[1]。
研究表明，在末次冰消期，云南洱海地区C3/C4植物的相对丰度存在显著的千年尺度波动，C4植物在16.0cal.ka B.P.和12.0cal.ka B.P.前后的两次扩张分别对应于Heinrich 1(H1) 和新仙女木(YD)事件。这些千年尺度波动与末次冰消期西南夏季风的变化具有很好的一致性，例如贵州董哥洞的石笋δ18O记录[58]、云南天才湖的孢粉[61]以及腾冲青海的沉积特征[63]等记录均反映了西南夏季风在H1和YD期间显著减弱，我国西南地区降水显著减少(图 3)。这表明在较为干旱的气候环境下，具有更高水分利用效率的C4植物更具有生长优势。此外，值得注意的是，C4植物开始快速减少的时期与西南夏季风明显增强的时间也基本一致，而显著滞后于冰芯记录的pCO2开始上升的时间。因此，在末次冰消期，西南夏季风的降水更可能是洱海地区C3/C4植物的相对丰度变化的主要环境因子，而pCO2的降低进一步促进了C4植物在末次冰盛期的扩张。
在全新世期间，云南洱海地区C3/C4植物的相对丰度变化趋势也与西南夏季风的演化相吻合。西南地区的石笋δ18O记录表明全新世以来夏季风逐渐减弱[58]；星云湖孢粉定量重建的近8ka降水逐渐减少，与石笋δ18O的变化趋势相一致[64]。虽然湖泊沉积记录中不同的代用指标的敏感度存在差异[65~67]，但青藏高原众多的湖泊记录均显示气候适宜期出现在早中全新世，晚全新世的有效湿度显著降低[68]。尽管该阶段pCO2呈现出自早全新世先逐渐下降，到晚全新世反转上升的变化趋势[41, 69]，但总体来说变幅较小，相对稳定，对植物组成的影响可以忽略。此外，在千年尺度上，洱海地区C4植物在8.2~7.1cal.ka B.P.和5.7~4.7cal.ka B.P.显著扩张(图 3)，分别对应于北大西洋的冷事件“Bond 5”和“Bond 4”[70]。研究表明，北大西洋的冰筏事件可能通过温盐环流影响低纬地区的气候，导致热带辐合带南移，西南夏季风减弱[58, 71]。然而，不同地区记录的弱夏季风事件的起讫时间、持续时间和变化幅度等方面存在较大的差异。例如，贵州董哥洞石笋δ18O记录显示夏季风在8.4~8.1ka显著减弱[58]；四川红原泥炭纤维素δ13C记录表明气候在8.6~8.1cal.ka B.P.期间比较干旱[71]；青海湖的多指标研究也表明了气候在8.4~8.0cal.ka B.P.期间显著恶化[72]；而青藏高原东南缘伍须海和那龙错的孢粉记录却显示冷事件“Bond 5”对应的干旱期持续时间较长，结束于7.0cal.ka B.P.左右[73, 74]。对于湖泊沉积物，事件前后缺少可靠的年代界定，特别是植物残体的AMS14C年代数据偏少，而全有机质测年材料往往受到湖盆“碳库效应”的影响[75]，可能是8.2ka事件发生区域差异的主要原因。
6 结论近20ka以来云南洱海沉积物δ13CEC值变化范围在-30.0 ‰ ~-20.8 ‰之间，平均约为-25.9 ‰，全新世δ13CEC值整体偏负于晚冰期。燃烧过程、温室气体浓度和气候变化产生的同位素分馏效应对δ13CEC影响较小，洱海沉积物的δ13CEC主要反映了末次冰盛期以来湖区C3/C4植物相对丰度的变化。在14.7cal.ka B.P.之前，洱海地区陆地植被为C3/C4植物混合类型，C4植物比较丰富的；此后C4植物相对丰度逐渐减少，在早全新世降到最低；在晚全新世C4植物有所扩张，但仍以C3植物为主导。洱海沉积物的δ13CEC反映的区域植被变化与低纬地区海洋、湖泊沉积物正构烷烃单体δ13C指示的C3/C4植物相对丰度变化基本一致，表明低纬地区末次冰期C4植物相对较多。此外云南洱海地区C4植物的相对丰度在15.7cal.ka B.P.、12.0cal.ka B.P.、8.0cal.ka B.P.以及5.0cal.ka B.P.前后出现了持续近千年的高值，这些变化与西南夏季风千年尺度的快速减弱事件密切相关，表明夏季风降水是云南洱海地区C3/C4植物演化的主要驱动因素。
致谢: 蒋庆丰、纪明、陈嵘、李艳玲、肖霞云等协助完成野外采样，在此深表感谢。本研究由国家重点研发计划项目(批准号：2016YFA0600502)、科技基础性工作专项项目(批准号：2014FY110400) 和国家自然科学基金项目(批准号：41572337) 共同资助。
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② College of Geography and Environment, Shandong Normal University, Jinan 250014;
③ University of Chinese Academy of Sciences, Beijing 100049)
Abstract
The variation of the relative abundance of C3/C4 plant in terrestrial vegetation since the Last Glacial Maximum(LGM)and its driving mechanism is a key research question in paleoecology. However, there is a lack of record based on the relative abundance of C3/C4 plant from southern part of China, particularly the high resolution paleoclimate records are sparse. Whether climate and/or atmospheric CO2 concentration are the most important driving factors for these changes remain to be elucidated. Black carbon or elemental carbon is produced by incomplete combustion of biomass and fossil fuels. Due to the inert characteristics of the elemental carbon, effect on the photochemical and microbial reactions after deposition can be negligible. Therefore black carbon can be well-preserved in the sediments for a long period of time. During the process of combustion, different isotopic fractionation effect exhibits as the chemical composition of biomass varies. However, the variation in the carbon isotopic composition(δ13CEC)is relatively small when compared to the plant δ13C. In the past few years, there has been increasing attention received on developing the method that uses δ13CEC to reconstruct vegetation and climate change from lacustrine sediments. In this study, a 684-cm sediment core(25°51.02'N, 100°11.02'E) was recovered at the depth of 13.5m of Erhai Lake(25°25'~26°16'N, 99°32'~100°27'E), Yunnan Province. The core was sub-sampled at 4cm intervals and a total of 171 carbon isotope(δ13CEC)samples were analyzed throughout the record. The relationship between the relative abundance of C3/C4 plant and the paleoenvironmental changes from Erhai Lake, Yunnan Province was investigated. The record covers periods since 19.4cal.ka B.P. to the present and is based on the analysis of elemental carbon δ13CEC and AMS radiocarbon(14C)dates using 13 plant remains and charcoal samples from the lake sediment. The results show that during the LGM, the composition of vegetation around Erhai Lake catchment contains a mix of C3 and C4 plants, and C4 plant is the dominant type during this period. The relative abundance of C4 plant decreased abruptly since 14.7cal.ka B.P. and has reached the minimum abundance in the Early Holocene. In the Late Holocene, the C4 plant shows an expansion to some degree however, the overall composition of vegetation is still dominated by C3 plant. This δ13CEC record from Erhai Lake is consistent with the relative abundance of C3/C4 plant recorded in other marine and lake sediments in the low latitudes. These records all together suggest that the abundance of C4 plant is relatively high and dominant the vegetation composition during the LGM. In addition, the C4 plant in the Erhai Lake catchment also shows significant expansion around 15.7cal.ka B.P., 12.0cal.ka B.P., 8.0cal.ka B.P. and 5.0cal.ka B.P. These abrupt changes were likely related to the rapid decline and weakening of the Indian Summer Monsoon on the millennium scale. In summary, this study concludes that the relative abundance of C3/C4 plant and the long-term vegetation evolution in Yunnan China were mainly driven by the variation in regional precipitation, which is controlled by the Indian Summer Monsoon since the LGM.