Regional characteristics and mechanisms of lake water level decline in the Tibetan Plateau since 2 000 years ago
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摘要:
基于亚洲夏季风与西风的影响范围将青藏高原划分为3个研究区，通过对比湖泊沉积物中多代用指标与晚全新世火山活动、北半球温度和亚洲季风指数，探讨了2 kaBP前后高原湖泊水位下降的原因，并分析了不同区域湖泊对气候波动响应的空间差异。结果表明，青藏高原西南部湖面水位下降幅度大于西北部，更甚于高原东北部。这可能是因印度夏季风（Indian Summer Monsoon，简称ISM）强度减弱，高原西南部的湖泊更依赖于ISM降水的补给，因此对该季风所带来的水汽通量的减少更加敏感。而且，该时期的北大西洋涛动（North Atlantic Oscillation，简称NAO）的位相由负转正，使得青藏高原北部水汽辐合增强、降水偏多而南部降水偏少，进而导致高原南部湖面水位下降幅度普遍大于北部湖泊。导致青藏高原气候趋于冷干的主要原因，本文归因于该阶段厄尔尼诺（EI Niño）的加强。除此之外，该时期南半球环状模（Southern Annular Mode，简称SAM）冬夏季的不同位相也通过复杂的海气耦合过程，跨越赤道对青藏高原气候起到了降温减湿的作用。
Abstract:The Tibetan Plateau (TP) was divided into three zones based on the influence of the Asian summer monsoon and the westerlies. By comparing multiple proxy indicators in sediments with late Holocene volcanic activity, the Northern Hemisphere temperatures, and the Asian monsoon index, the reasons for the decline in plateau lake levels ～2 kaBP were explored and the spatial differences in lake responses to climate fluctuations in the different zones were analyzed. Results show that the decline in lake water level in the southwestern part of the TP is greater than in the northwestern part, and even greater in the northeastern TP. This may be due to the weakening in the intensity of the Indian Summer Monsoon (ISM), which made lakes in the southwestern TP more dependent on the ISM precipitation replenishment and thus more sensitive to the reduction in water vapor flux brought by the ISM. Moreover, during this period, the phase of the North Atlantic Oscillation (NAO) shifted from negative to positive, leading to the increase in water vapor convergence in the northern part of the TP with more precipitation there, while the southern part of the TP received less rainfall, resulting in a generally greater decline in water levels in the southern lakes compared to those in the north. The main cause of the climate turning to colder and drier in the TP ～2 kaBP is attributed to the intensification of El Niño. In addition, the different phases of the Southern Annular Mode in winter and summer through complex ocean-atmosphere coupling processes crossing the equator, also played a role in cooling and dehumidifying the climate of the TP.
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Keywords:
- climate change /
- Tibetan Plateau lakes /
- Indian Summer Monsoon /
- ENSO /
- NAO
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青藏高原作为地球上海拔最高的高原，平均海拔超过4000 m，面积约为250×104 km²，被誉为“世界屋脊”和“地球第三极”[1]。其位于东亚季风、印度季风以及盛行西风的交汇影响区（图1），使得高原上的湖泊对气候变化响应极为敏感[2]。龙木错[3]、青海湖[4]、色林错[5]、纳木错[6-7]、当惹雍错[8]、扎日南木错[9-10]、达则错[11]、当穹错[12]等湖泊的湖面水位波动曲线为研究全新世青藏高原湖泊演化提供了有价值的信息。气候变化是全新世时期湖泊扩张与收缩的主要因素，青藏高原上的大部分湖泊对末次冰消期以来气候环境变化响应的趋势总体上相似。末次冰消期以来，青藏高原的气候出现了多次突变，并且呈现出快速降温、缓慢升温的现象[13]。在全新世早中期，青藏高原气候整体呈相对温暖湿润的状态，但在中全新世后，气候逐渐向冷干转变，植被开始逐渐退化[14-15]。晚全新世时期，湖泊水位出现明显的下降。
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图 1 青藏高原地形图以及文中提及的气候代用指标记录的分布位置本研究将其划分为3个研究区域，其中Ⅰ区包含：1.郭扎错[18]，2.龙木错[3]，3.班公错[19]，4.阿翁错[20]，5.Tso Moriri湖[21]；Ⅱ区包含：6.哈拉湖[22]，7.青海湖[4]，8.黑海[23]，9.苦海[24]，10.班德湖[25]，11.Nir'pa Co[26]，12.天才湖[27]；Ⅲ区包含：13.扎日南木错[9-10]，14.当惹雍错[8]，15.当穹错[12]，16.达则错[11]，17.色林错[5]，18.班戈错[28]，19.纳木错[7]，20.兹格塘错[29]。粉色虚线为青藏高原现代亚洲季风区与非季风区分界线[30]，黑色实线为青藏高原边界[31]，使用ArcGIS Pro3.0.2（www.esri.com）以世界地形基准图为底图（Esri、TomTom、Garmin、FAO、USGS、NOAA）创建地图。Figure 1. Topographic maps of the Tibetan Plateau and the locations of proxy climate indicator records mentioned in the textThe three study zones were delineated in this study, among which Zone I includes: 1. Guozha Co[18], 2. Longmu Co[3], 3. Bangong Co[19], 4. Aweng Co[20], 5. Tso Moriri lake[21]; Zone Ⅱ includes: 6. Hala Lake[22], 7. Qinghai Lake[4], 8. Heihai Lake[23], 9. Kuhai Lake[24], 10. Lake Bande[25], 11. Nir'pa Co[26], 12. Tiancai Lake[27]; Zone Ⅲ includes: 13. Zhari Nam Co[9-10], 14. Tangra Yumco[8], 15. Tanqung Co[12], 16. Dagze Co[11], 17. Seling Co[5], 18. Bange Co[28], 19. Nam Co[7], 20. Zigetang Co[29]. The pink dashed line segments represent the modern boundary between the Asian monsoon and non-monsoon zones on the Tibetan Plateau[30]. The black line indicates the boundary of the Tibetan Plateau[31]. Maps were created using ArcGIS Pro3.0.2 (www.esri.com) with the World Terrain Basemap (Esri, TomTom, Garmin, FAO, USGS, NOAA).根据米兰科维奇理论，北半球夏季中高纬地区太阳辐射强度受地球轨道参数变化的影响，而这又与厄尔尼诺-南方涛动（ENSO）的振幅密切相关[16-17]。距今两千年前（2 kaBP）太阳辐射强度的季节性变化显著放大了ENSO的循环[32]，减弱了当时的印度夏季风强度，这无疑成为了青藏高原两千年前气候变冷变干的证据之一。同一时期的大气系统并非孤立地影响区域气候，而是通过一系列综合且复杂的海气耦合过程，共同作用于高原上空的西风环流及季风环流，从而导致降水的变化。因此，本文将分别探讨北大西洋涛动（NAO）、太平洋涛动（PDO）以及南半球环状模（SAM）等对青藏高原气候波动的影响，旨在解释青藏高原最近两千年湖面水位下降与气候冷干的联系，并揭示造成气候变化的潜在驱动力。
青藏高原上不同区域湖泊对水源补给丰缺的响应存在着差异[33]。因此，为探究不同湖泊湖面水位变化的时空差异，本文根据最近学者对不同大气环流在高原上的影响范围[30-34]，将青藏高原划分为Ⅰ、Ⅱ和Ⅲ三个区域（图1）。Ⅰ区主要包含了受盛行西风带降水影响的青藏高原西北部区域。该区海拔较高，气候极端干旱，区域内湖泊沉积物很好地记录了西风环流的变化历史。该区域内研究较多的湖泊有班公错、龙木错和郭扎错等。然而，在全新世季风最强时期，印度夏季风也为本区带来了一定量的降水[3, 35-36]。Ⅱ区是主要受东亚夏季风和印度季风共同补给的高原东部区域。前人对该区内青海湖、苦海、哈拉湖等研究较多，这些湖泊对东亚季风和印度季风降水变化极为敏感。Ⅲ区包含了主要由印度夏季风降水（ISM）所补给的高原西南部湖泊，如当惹雍错、当穹错、色林错、纳木错、达则错等，此区较其他两个区域更加接近ISM的水汽来源。由于湖泊沉积物中碳酸盐物质的同位素组成对降水、蒸发及水汽输送等区域大气过程有较强的指示意义[37]，因此本文收集了前人在这3个研究区内已报道的湖相碳酸盐沉积物δ18O数据及其他代用指标、不同湖泊的湖面水位变化曲线及其对应的年代数据、火山活动记录、亚洲季风指数、北半球中高纬温度记录等，探析青藏高原湖泊2 kaBP前后的演化历史，了解其对气候的响应及出现区域差异的原因。
1. 研究方法
1.1 青藏高原湖泊水位变化记录
古湖岸线提供了过去湖泊水位变化的证据，因此通过测定其年代可以重建湖面波动历史，并推断出流域过去的湿度变化[38]。目前高原上有许多的古湖岸线被测年和报道，也有较多的基于古湖岸线开展的全新世湖泊水位变化的研究。本文收集了10条前人所重建的青藏高原湖面变化曲线，其中青海湖[4]、色林错[5]、达则错[11]、扎日南木错[9-10]、班戈错[28]、纳木错[6-7]的湖面变化由古湖岸线的光释光年龄重建；当穹错[12]和当惹雍错[8]的湖面水位变化重建基于其湖滨沉积物光释光年龄[39]和宇生核素年龄[40]；龙木错[3]根据其古湖岸线沉积物与近岸湖相沉积物的光释光年龄、宇生核素年龄以及U/Th年龄重建了湖面波动曲线；苦海[24]则结合沉积物中碳酸盐、矿物相、钛通量、粒度和介形虫丰度等分析，重建了全新世湖面变化。通过观察发现在晚全新世约2 kaBP，高原上湖泊集体开始萎缩，湖面水位呈现出统一较大幅度下降的趋势（图2）。
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图 2 全新世青藏高原的湖面变化Figure 2. The Holocene lake levels changes in the Tibetan Plateau1.2 Ⅰ区的古气候代用指标记录
Ⅰ区位于青藏高原西北部，地势高亢，气候干寒，有大量冰川发育，属高原亚寒带干旱区[41]。该区主要受西风带的影响[35]，湖泊的补给多来源于冰川融水以及非季风降水。本文收集的代用指标来自于区内冰川以及4个湖泊的已发表记录（图3）。其中，降水的δ18O数据选用了崇测冰芯[42]中的记录。降水δ18O对气温变化具有高敏感性并且有较高的时间分辨率。其基本原理是水汽海陆循环中δ18O是地表温度的函数[43]。水的氧同位素组成包含了H216O和H218O，而其中“重水”H218O含量少。与H216O相比，H218O难于蒸发却易于凝结[44]，因此地表温度的变化驱动着水汽中δ18O含量。
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图 3 Ⅰ区过去11 ka青藏高原古气候记录对比a: 阿翁错自生碳酸盐δ18O（黄色实线）和介形虫壳体δ18O（墨绿色点）记录[20]，b: 崇测冰芯降水δ18O记录[42]，c: Tso Moriri湖自生碳酸盐δ18O记录[21]，d: 班公错自生碳酸盐δ18O记录[19]，e: 郭扎错自生碳酸盐δ18O记录[18]，f: 郭扎错沉积物磁化率记录[18]，g: 郭扎错沉积物LOI记录[18]。Figure 3. Comparison of paleoclimate records in Zone I from the past 11 ka in the Tibetan Plateaua: Records of autogenic carbonate δ18O (yellow line) and ostracods shell δ18O (dark green dots) in Aweng Co[20], b: Precipitation δ18O records from the Chongce ice cores[42], c: Autogenic carbonate δ18O records from Tso Moriri lake[21], d: Autogenic carbonate δ18O records from Bangong Co[19], e: Autogenic carbonate δ18O records from Guozha Co[18], f: Magnetic susceptibility records from Guozha Co sediments[18], g: Loss on ignition (LOI) records from Guozha Co sediments[18].在阿翁错[20]、Tso Moriri湖[21]、班公错[19]以及郭扎错[18]的湖相沉积记录中都收集了各自湖相沉积自生碳酸盐的δ18O数据，其可以很好地反映湖水补给来源的丰缺情况，即降水或冰川的进退过程。除此之外，还选用了郭扎错沉积物的磁化率（MS）与烧失量（LOI）作为参考。MS与湖泊沉积物中磁性矿物的含量直接相关，表明湖泊集水区径流对碎屑物质的侵蚀和运输能力；LOI反映了湖泊内的初级生产力，当冰川融水带来的碎屑物质增多时，会降低湖泊沉积物中的有机质含量，即LOI为低值。
1.3 Ⅱ区的古气候代用指标记录
Ⅱ区主要包含青藏高原东部东亚夏季风与印度夏季风交互影响的过渡区域。在该区选择的8个湖相沉积物以及相关记录都很好地对应了晚全新世东亚夏季风（EASM）的演化趋势（图4）。其中自生碳酸盐的δ18O数据来自于哈拉湖[22]、苦海[24]以及班德湖[25]；德令哈具有精确年代控制的高分辨率树轮δ18O年表[45]记录了亚洲夏季风降水的高低频变化；青海湖则选用了全新世中晚期以来的花粉数据[46]和基于$\rm U^K_{37} $重建的湖区温度记录[47]；黑海使用了沉积物中由介形类反映的湖水电导率[23]作为气候变化的参考；Nir’pa Co沉积物中的粉砂含量比反映了地表径流搬运碎屑能力的强弱[26]；天才湖则选用了由brGDGT推断出的湖区年均温重建记录[27]。
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图 4 Ⅱ区过去11 ka青藏高原古气候记录对比a: 哈拉湖自生碳酸盐δ18O记录[22]，b: 德令哈树轮δ18O序列[45]，c: 青海湖木本植物花粉（黄色曲线）、草本植物花粉（绿色曲线）以及总花粉序列（蓝色曲线）[46]，d: 黑海湖基于介形类重建的湖水电导率[23]，e: 苦海湖自生碳酸盐δ18O记录[24]，f: 班德湖自生碳酸盐δ18O记录[25]，g: Nir’pa Co沉积物粉砂含量记录[26]，h: 青海湖温度重建序列[47]，i: 天才湖温度重建序列[27]。Figure 4. Comparison of paleoclimate records in Zone II from the past 11 ka in the Tibetan Plateaua: Autogenic carbonate δ18O records from Hala Lake[22], b: Tree ring δ18O sequences from Delingha[45], c: Pollen sequences from arboreal plants (yellow curve), herbaceous plants (green curve), and total pollen (blue curve) in Qinghai Lake[46], d: Lake water conductivity reconstructions based on ostracods from Heihai Lake[23], e: Autogenic carbonate δ18O records from Kuhai Lake[24], f: Autogenic carbonate δ18O records from Lake Bande[25], g: Silt content records from Nir’pa Co sediments[26], h: Temperature reconstruction sequences from Qinghai Lake[47], i: Temperature reconstruction sequences from Tiancai Lake[27].1.4 Ⅲ区的古气候代用指标记录
Ⅲ区位于青藏高原西南部，气候以半干旱为主，区内绝大部分湖泊流域地处ISM降水控制地带。因此我们在Ⅲ区范围内选择的3个湖泊沉积物中的代用指标都指示了ISM的强弱（图5）。纳木错沉积物中的K[48]和Ti含量[49]直接反映了流入湖中的成矿成分以及季风降水的强度与频率[50]；而Ca含量[49]则指示了湖泊碳酸盐的累积过程以及湖面变化。在色林错选用了Ti含量[51]与自生碳酸盐δ18O记录[52]作参考用来推断湖区干湿变化。前人利用兹格塘错[29]的孢粉记录重建了该湖区植被类型的演替过程，间接表达了气候的干湿变化。
2. 结果与讨论
2.1 2 kaBP青藏高原湖面下降的空间差异
全新世期间青藏高原湖面变化曲线表明，最近2 000年青藏高原上的湖泊水位普遍下降（湖泊范围萎缩），且下降速度较全新世其他时期更快，下降幅度更大（图2）。从分区来看，最近2 000年高原西南部Ⅲ区湖泊水位下降幅度大于西北部Ⅰ区，更甚于东部Ⅱ区。位于Ⅲ区的班戈错、达则错、纳木错、色林错、扎日南木错、当惹雍错和当穹错的湖面自约2 kaBP以来，分别下降了约23、30、34、40、44、70和100 m；西北部Ⅰ区的龙木错下降了约25 m；而东部Ⅱ区的青海湖水位降低了8 m，苦海与1.5 kaBP的低湖面水位也仅相差约10 m。
现代气候背景条件下，青藏高原大气环流的特点是南部受ISM控制，北部和西部受西风影响[53]，EASM和ISM都对东部的气候有影响[54]。由于北半球太阳辐射的变化，季风与西风带的控制范围并非一成不变，这同时造成了青藏高原不同地区温度与降水的波动，进而影响代用指标对气候变化的响应和记录。Ⅰ区的阿翁错位于高原内现代西风带的南缘，其沉积物中自生碳酸盐δ18O在2 kaBP前后表现出频繁的波动（图3a）。这可能说明因太阳活动引起的外强迫导致西风带发生了移动。而居于高原西北缘的崇测冰芯的降水δ18O则在该时段呈现明显的下降，表明Ⅰ区气候向冷干变化，这与班公错以及郭扎错代用指标所表现的气候波动趋势相同。Tso Moriri湖的自生碳酸盐δ18O也在约2 kaBP前后呈现高值，说明高原西北部喜马拉雅南麓降水减少。Ⅱ区的气候指标按纬度位置观察，发现在高原东部越接近南部的湖泊，其自生碳酸盐δ18O波动频率越小，即呈现明显连续的高值。哈拉湖δ18O在约3 kaBP后稳定降低，但在约2 kaBP开始出现剧烈波动。与之相反的是纬度更低一些的苦海和班德湖，二者自生碳酸盐δ18O都在2～1.5 kaBP出现稳定上升，这与青海湖和天才湖的温度重建结果的变化相一致，即气候在该时间范围内转入相对冷干阶段。青海湖的花粉数据显示，湖区花粉浓度有所增长，但主要是以草本孢粉浓度增长为特征，木本花粉数量急剧下降，似乎显示出该期以冷干的气候为特征[46]。黑海基于介形类重建的湖水电导率也显示出蒸发加剧径流减少的趋势[23]。Nir’pa Co的粉砂含量的降低，也表明当时流域内水流量减少，地表径流搬运能力较弱，气候偏干旱。Ⅰ区与Ⅱ区的研究结果共同证明了在～2 kaBP的时候，西风带的影响范围可能扩大并且西风的强度有所增加，使得区域内湖泊接受的季风性降水减少，而非季风性降水可能增加。而Ⅰ区的湖泊水位下降幅度大于Ⅱ区的原因可能与其所处的海拔高度有关。高原西北部湖泊的流域海拔较高且大部分湖区内冰川广布，由于冰雪对太阳活动和火山活动等负强迫存在着强烈的正反馈机制[55]，导致Ⅰ区湖泊对区域气候变冷的响应要敏感于海拔较低处的湖泊。此外，造成高原西北部高海拔地区降水的天气系统多为对流层西风带上层天气尺度的低压系统，因此该地区的降水机制也区别于低海拔地区[56-57]。
Ⅲ区3个湖泊的不同代用指标在图5中都统一指示了气候向冷干转变的特征。纳木错沉积物中K、Ca、Ti元素的含量降低反映了ISM降水强度的减弱；色林错的Ti以及自生碳酸盐δ18O也反映出湖区强的风沙活动或降低的湿度；对兹格塘错流域植被类型的重建显示出～2 kaBP流域内以高寒草原为典型植被类型的特征。造成上述现象的主要原因可能是因为该时段ISM强度的减弱以及季风水汽通量降低。Ⅲ区湖泊水位下降幅度大于Ⅰ、Ⅱ区的原因可能是该区的湖泊更接近ISM的水汽源头，来自印度洋的水汽首先到达高原南部并形成降水，其降水强度向北部递减[11, 58]，因此比其他区域的湖泊更加依赖于ISM相关的降水。另一个原因可能是～2 kaBP，大西洋多年代际振荡（AMO）指数变化的总趋势持续上升[59]。有研究认为，AMO的暖位相与增强的中纬度西风带有关[60]。西风带导致北大西洋上空的负地面气压异常延伸至中纬度欧亚大陆[61]。这些因素反过来又削弱了西伯利亚-蒙古高压系统，并减弱了东亚冬季风的强度[62]。因此冬季湖泊净蒸发量的减少抵消了一部分因夏季降水强度减弱带来的水位下降。并且因西风增强所带来的非季风性降水的增多也可能在一定程度上弥补了季风性降水减少带来的亏损。Ⅰ区的龙木错和Ⅱ区的青海湖都有可能在约2 kaBP受到AMO位相变化所带来的高原气压场异常的影响。与AMO有关的另一种大西洋海气耦合模态是NAO。该时期NAO位相的百年尺度变化趋势由负转正（图8a）。在NAO指数较高的年份，青藏高原北部水汽输送通量强度增加，且水汽辐合增强（图7）。与之相对应的是，高原南部水汽输送通量强度减弱。在这种水汽输送和环流模式配置下，可能使得高原北部降水偏多而南部降水偏少[63]，进而导致高原南部湖面下降幅度普遍大于北部湖泊。
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图 8 类NAO、类SAM以及类PDO条件百年尺度的指数变化Figure 8. Changes in centennial-scale indices of the NAO-like, SAM-like, and PDO-like scenarios![]()
图 7 NAO、IOD以及ENSO对青藏高原的湿度影响示意图Figure 7. Schematic diagram of the impacts of NAO, IOD, and ENSO on the humidity of the Tibetan Plateau2.2 青藏高原湖面下降的潜在驱动力
2.2.1 赤道印度洋与太平洋地区海气耦合系统的演化记录
青藏高原上的湖泊水位波动在全新世早期受到多种因素的影响，尽管有些湖泊的水位变化主要由冰川融水控制，但大部分湖泊最高水位的出现与亚洲夏季风强度在10～8 kaBP的增强相吻合，表明多数湖泊水位主要受季风控制[64]。前人根据青藏高原湖泊沉积物记录推断出的水文气候变化通常与用花粉指数重建的地表气温推断出的变化相似[29]，说明在2.1～1.3 kaBP之间ISM降雨量减少与气温下降同时发生。青藏高原湖泊沉积物中所记录的印度夏季风演化趋势具有一致性，揭示了约2 kaBP使气候变冷变干的大规模强迫机制。
ENSO是热带太平洋大尺度海气相互作用的产物，是全球海洋-大气耦合系统在年际尺度上最主要的模态[65]。印度夏季风与印度-太平洋海面温度以及类ENSO条件的百年和千年尺度的海洋-大气变化密切相关，现代ENSO-ISM的关系已经被建立[66-68]。青藏高原湖泊沉积物的指标分析结果与其他地区古气候记录对比后发现，在3.3 ～2.2 kaBP之间赤道东太平洋变冷，并且印度-太平洋变暖，表明当时为类La Niña条件，这与高原气候在该时期的相对暖湿相吻合。来自加拉帕戈斯群岛El Junco火山口湖的El Niño记录为晚全新世ENSO变率提供了可靠的证据（图6a），该变率与印度太平洋海温记录一致[69]。在3.3～2.4 kaBP的El Junco记录中，砂量的低值通常被解释为对湖泊流域降水减少的响应。这样的情况会在La Niña年发生，间接表明青藏高原在这个时期的气候条件更加湿润温暖。El Junco在2.4～1.3 kaBP之间出现了峰值砂量，表明该时期持续的El Niño条件。在该阶段，印度太平洋海温出现负异常，进而导致ISM降水减少。这与高原湖泊水位下降、湖泊沉积物指标反映出的气候干旱相一致。
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图 6 11 kaBP以来大尺度海气耦合系统指数与北半球古降水/水分变化以及其他记录的比较a: 加拉帕戈斯群岛El Junco湖的含砂量记录，可视为赤道太平洋东部厄尔尼诺活动的代表[69]；b: 智利南部Lago CipresesLC: 湖泊记录中Pilgerodendron uviferum丰度的百分比，可解释为类SAM条件负相位条件的代表[84]；c: 火山影响活动[85]；d: Cariaco盆地沉积物中的Ti记录[78]；e: 北半球中高纬度温度变化重建记录[83]，f: 亚洲中部干旱区湿润指数[54]；g: 冰川活动记录[54]，h: 纳木错Kclr记录[48]；i: 两条董哥洞δ18O记录[80-81]。Figure 6. Comparison between large-scale ocean-atmosphere coupling system indices and the Northern Hemisphere paleoprecipitation / moisture changes and other records since 11 kaBPa: Sediment sand content records from El Junco Lake in the Galápagos Islands, representing eastern equatorial Pacific El Niño activity[69, b: Percentage of Pilgerodendron uviferum abundance in lake records from, interpreted as a representation of SAM-like negative phase conditions[84]; c: Volcanic influence activity[85]; d: Ti records in sediments from the Cariaco Basin[78], e: Reconstruction records of temperature changes in the mid to high latitudes of the Northern Hemisphere[83]; f: Humidity index for the arid regions of central Asia[54]; g: Glacier activity records[54]; h: Nam Co Kclr records[48];i: Two carbonate δ18O records from Dongge Cave[80-81].ENSO的历史与地轴的进动强迫密切相关[16-17]，日照季节性的变化明显放大了约5 kaBP和2 kaBP的ENSO循环[32]。前人通过展示了覆盖过去12 000年受ENSO变率强烈影响的厄瓜多尔南部拉古纳帕尔卡科查湖的沉积记录，将ENSO变率的长期趋势归因于轨道引起的太阳日照辐射的季节性变化，并发现ENSO振幅峰值出现在2～1 kaBP左右[70]。
赤道印度洋上与ENSO息息相关的另一种海气相互作用现象称为印度洋偶极子（IOD）[71]。来自塞舌尔的珊瑚记录[72]和印度洋东部珊瑚的全新世记录表明两千年前IOD的活动增强[73]。其正位相使得东部印度洋海温处于低值状态，间接引起亚洲夏季风强度减弱。通过近千年的地质记录和IOD变率与ENSO变率的对比，发现二者变率具有较好的一致性。IOD与ENSO之间呈现出较强的非线性关系，且大部分强的正IOD事件对应着强El Niño事件[74]。强El Niño位相下，IOD事件相关的海温与风场表现出很强的响应，强于La Niña事件产生的响应，表现为强的非对称性[75-76]。因此，ENSO可通过激发非对称的大气遥相关对IOD强度产生调制作用，印度洋海表温度偏度很大程度上是由强El Niño事件导致的强正IOD事件所贡献的[77]（图7）。因此本研究虽然将约2 kaBP青藏高原气候变冷变干的现象主要归因于ENSO循环的放大，但ISM降水的减少并非由于ENSO的单一作用。
上述两类海气耦合系统所导致的赤道印度洋与太平洋海温异常可能会引起热带辐合带（ITCZ）的平均纬度的变化。来自委内瑞拉海岸附近Cariaco盆地的Ti浓度数据可用于推断ITCZ的移动（图6d）[78]。北半球ITCZ的南移通常与亚洲夏季季风的弱化相关联，这种南移导致陆地上空形成高气压区，从而无法吸引来自海洋的湿润空气，减少了降水量[79]。纳木错元素K含量[48]（图6h）以及董哥洞自生碳酸盐δ18O[80-81]（图6i）记录的亚洲夏季风减弱与ITCZ的位置变化有较好的一致性（图6c）。另一方面，ITCZ的位置变化也会对北半球的温度产生影响。由于ITCZ是热带降水的主要区域，其位置的改变会影响热量和水分的分布[82]，进而影响温度模式。因此，ITCZ的北（南）移可能导致北半球一些地区温度升高（降低）。图6d所表现出来的ITCZ在2 kaBP的南移趋势与北半球中高纬度温度变化重建记录[83]（图6e）和亚洲中部干旱区湿润指数[54]（图6f）的下降相吻合。高原上的冰碛物年龄密度分布也显示出该阶段冰川处于稳定的扩张状态[54]（图6g），火山活动影响指数[85]在该时期也出现了一次小的峰值（图6c），表现出气候相对寒冷。
2.2.2 热带外地区大气模态对青藏高原气候的影响
SAM是南半球中高纬度大气环流变率的主导模态[86]，它的气候效应不仅仅作用于南半球，还可以向北延伸越过赤道达到北半球。海气耦合过程在SAM对北半球气候的影响中扮演着重要的角色，与SAM信号的跨季节存储和由南半球向北半球的传播均有密切关系[87]。前人的研究发现ENSO和SAM的相互作用对日照强迫存在千年尺度的响应[32]，二者变率的负相关关系在约5.8 kaBP之后建立，并在4.6 kaBP达到峰值，表明热带和热带外大气系统变率模式的连贯变化[84]（图6a , 6b）。在2 kaBP左右，ENSO的强正位相对应了SAM的高负位相。具体到热带太平洋地区，12—2月SAM正位相对应次年3—5月赤道太平洋信风加强，热带中太平洋海温偏冷，这一冷海温和信风的增强进一步通过正反馈机制将信号放大，反之则相反[88]。因此，当12—2月SAM处于负位相时，其导致的热带太平洋暖海温异常叠加在ENSO活动对应的海温异常上，使得3—5月ENSO暖事件偏强，而冷事件偏弱[89]。前人通过IPSL-CM5A2模拟的SAM季节性变化指数发现，2 kaBP时12月份的SAM一直处于负相位[90]。其造成的海温异常信号通过海气耦合过程影响了次年春季El Niño的强度，间接减少了季风在该时期给青藏高原带来的水汽通量。
此外，高原积雪通过其反照率和导热性调控高原的热力作用，即高原积雪增多时，高原地表反照率增强，地表温度降低，减弱地表向大气的感热和潜热输送，进而削减高原的热力作用[91-92]。在北半球夏季，虽然在高原东部大部分地区的积雪都已融化，但在高原西部和南边缘一侧高海拔地区仍然有积雪存在。这些积雪一部分来源于前期冬春季的滞留，一部分来自其他外部强迫因子的影响。前人通过研究发现5月SAM正位相往往伴随着高原西部夏季积雪的异常增多[88]。二者之间复杂的海气耦合机制可简要概括为以下4个过程：首先，SAM正位相通过动力和热力作用，引起南半球中高纬度海温异常，具体表现为绕极海温的冷却以及中纬度的海温增暖[93-94]。其次，印度洋对这种海温异常的响应表现冷（～20°S）暖（～40°S）冷（～60°S）的印度洋经向三极子（IOT）海温模态[95]。再次，正IOT海温通常会导致对流层低层西印度洋出现高压反气旋，而东印度洋海洋性大陆低层则出现气旋式复合运动，从而激发出热带印度洋的纬向降水偶极子模态[88]。热带降水强对流活动中释放的凝结潜热，产生异常涡源，会在热带外地区激发低频行星尺度波扰动，通过遥相关引起热带外地区的气候异常[96-99]。最后，根据波射线理论[100]，该降水偶极子引起的潜热释放所产生的波源能够激发定常波，在中低层背景流的作用下传播到青藏高原地区，引起高原西部的气旋性上升环流，为积雪的生成和维持提供有力的动力条件[88]。该物理过程中， IOT海温和热带偶极子降水分别扮演了重要的“海洋桥”和“大气桥”作用。约2 kaBP时，SAM在5月份持续处于正位相（图8b），与之相关的空气温度在高原上空表现为负异常[101]，导致积雪不易融化，从而有利于高原西部地表低温状态的维持。
除此之外，前人的研究发现PDO很大程度地驱动着中国的降水[102]，其通过影响西伯利亚高压系统来控制东亚冬季和夏季风的强度[103]，而ENSO又可以在季节到年际时间尺度上对亚洲季风降雨产生直接影响。此外，该降水变化又与NAO密切相关[104]。这3种大气振荡呈现不同相位组合时，高原的水汽变化也会出现空间差异。王江林等使用从亚洲季风干旱集和全球帕默尔干旱指数（PDSI）数据集中提取的PDSI序列[107]，研究了整个青藏高原的水分空间和时间变化，发现当NAO和ENSO处于正相、PDO处于负相时（图8c），高原中南部极有可能出现干旱（图9）。约2 kaBP的3种振荡与上述的位相组合相吻合，因气候变干导致的湖泊净蒸发量激增可能也成为当时Ⅲ区湖泊水位下降极为剧烈的原因之一。
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3. 结论
本文根据盛行西风和亚洲夏季风的影响区域将青藏高原划分为3个研究区，并结合3个区域中代表性湖泊的水位变化曲线和沉积物中的多气候指标数据，分析了各区2 kaBP前后的气候波动机制以及湖面下降的空间差异。结果显示：
（1）青藏高原西南部湖面下降幅度大于西北部，更甚于东北部。3个研究区的代用指标都反映出了该阶段季风降水的减弱，但造成水位下降空间差异的原因可能是由于高原西南部Ⅲ区的湖泊更靠近ISM的水汽来源，因此较其他区域也更加依赖ISM的降水，对该季风所带来的水汽通量的减少更加敏感。并且该时期的NAO百年尺度的指数由负转正，引起青藏高原北部水汽辐合增强，降水偏多，而南部降水偏少，进而导致高原南部湖面下降幅度普遍大于北部湖泊。
（2）2 kaBP前后地球轨道参数的变化放大了ENSO循环，这直接影响了亚洲夏季风的强度。因此厄尔尼诺事件的加剧也成为了本研究认为导致约2 kaBP青藏高原气候变冷变干的主因，并且当时SAM冬季的负位相增强了次年春季的厄尔尼诺，间接对高原起到减湿作用；其5月的正位相又通过一系列的海气耦合过程维持了高原西部的夏季积雪，增强了地表反射率，起到了对高原的降温作用。海气耦合系统相互作用所造成的海温异常也通过影响北半球的温度和湿度，直接或间接地造成了当时高原气候的冷干条件。
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图 1 青藏高原地形图以及文中提及的气候代用指标记录的分布位置
本研究将其划分为3个研究区域，其中Ⅰ区包含：1.郭扎错[18]，2.龙木错[3]，3.班公错[19]，4.阿翁错[20]，5.Tso Moriri湖[21]；Ⅱ区包含：6.哈拉湖[22]，7.青海湖[4]，8.黑海[23]，9.苦海[24]，10.班德湖[25]，11.Nir'pa Co[26]，12.天才湖[27]；Ⅲ区包含：13.扎日南木错[9-10]，14.当惹雍错[8]，15.当穹错[12]，16.达则错[11]，17.色林错[5]，18.班戈错[28]，19.纳木错[7]，20.兹格塘错[29]。粉色虚线为青藏高原现代亚洲季风区与非季风区分界线[30]，黑色实线为青藏高原边界[31]，使用ArcGIS Pro3.0.2（www.esri.com）以世界地形基准图为底图（Esri、TomTom、Garmin、FAO、USGS、NOAA）创建地图。
Figure 1. Topographic maps of the Tibetan Plateau and the locations of proxy climate indicator records mentioned in the text
The three study zones were delineated in this study, among which Zone I includes: 1. Guozha Co[18], 2. Longmu Co[3], 3. Bangong Co[19], 4. Aweng Co[20], 5. Tso Moriri lake[21]; Zone Ⅱ includes: 6. Hala Lake[22], 7. Qinghai Lake[4], 8. Heihai Lake[23], 9. Kuhai Lake[24], 10. Lake Bande[25], 11. Nir'pa Co[26], 12. Tiancai Lake[27]; Zone Ⅲ includes: 13. Zhari Nam Co[9-10], 14. Tangra Yumco[8], 15. Tanqung Co[12], 16. Dagze Co[11], 17. Seling Co[5], 18. Bange Co[28], 19. Nam Co[7], 20. Zigetang Co[29]. The pink dashed line segments represent the modern boundary between the Asian monsoon and non-monsoon zones on the Tibetan Plateau[30]. The black line indicates the boundary of the Tibetan Plateau[31]. Maps were created using ArcGIS Pro3.0.2 (www.esri.com) with the World Terrain Basemap (Esri, TomTom, Garmin, FAO, USGS, NOAA).
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图 2 全新世青藏高原的湖面变化
Ⅰ区：龙木错[3]，Ⅱ区：青海湖[4]，苦海[24]，Ⅲ区：当穹错[12]，当惹雍错[8]，扎日南木错[9-10]，色林错[5]，班戈错[28]，纳木错[7]，达则错[11]。
Figure 2. The Holocene lake levels changes in the Tibetan Plateau
Zone I: Longmu Co[3], Zone Ⅱ: Qinghai lake[4], Kuhai lake[24], Zone Ⅲ: Tanqung Co[12], Tangra Yumco[8], Zhari Nam Co[9-10], Seling Co[5], Bange Co[28], Nam Co[7], Dagze Co[11].
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图 3 Ⅰ区过去11 ka青藏高原古气候记录对比
a: 阿翁错自生碳酸盐δ18O（黄色实线）和介形虫壳体δ18O（墨绿色点）记录[20]，b: 崇测冰芯降水δ18O记录[42]，c: Tso Moriri湖自生碳酸盐δ18O记录[21]，d: 班公错自生碳酸盐δ18O记录[19]，e: 郭扎错自生碳酸盐δ18O记录[18]，f: 郭扎错沉积物磁化率记录[18]，g: 郭扎错沉积物LOI记录[18]。
Figure 3. Comparison of paleoclimate records in Zone I from the past 11 ka in the Tibetan Plateau
a: Records of autogenic carbonate δ18O (yellow line) and ostracods shell δ18O (dark green dots) in Aweng Co[20], b: Precipitation δ18O records from the Chongce ice cores[42], c: Autogenic carbonate δ18O records from Tso Moriri lake[21], d: Autogenic carbonate δ18O records from Bangong Co[19], e: Autogenic carbonate δ18O records from Guozha Co[18], f: Magnetic susceptibility records from Guozha Co sediments[18], g: Loss on ignition (LOI) records from Guozha Co sediments[18].
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图 4 Ⅱ区过去11 ka青藏高原古气候记录对比
a: 哈拉湖自生碳酸盐δ18O记录[22]，b: 德令哈树轮δ18O序列[45]，c: 青海湖木本植物花粉（黄色曲线）、草本植物花粉（绿色曲线）以及总花粉序列（蓝色曲线）[46]，d: 黑海湖基于介形类重建的湖水电导率[23]，e: 苦海湖自生碳酸盐δ18O记录[24]，f: 班德湖自生碳酸盐δ18O记录[25]，g: Nir’pa Co沉积物粉砂含量记录[26]，h: 青海湖温度重建序列[47]，i: 天才湖温度重建序列[27]。
Figure 4. Comparison of paleoclimate records in Zone II from the past 11 ka in the Tibetan Plateau
a: Autogenic carbonate δ18O records from Hala Lake[22], b: Tree ring δ18O sequences from Delingha[45], c: Pollen sequences from arboreal plants (yellow curve), herbaceous plants (green curve), and total pollen (blue curve) in Qinghai Lake[46], d: Lake water conductivity reconstructions based on ostracods from Heihai Lake[23], e: Autogenic carbonate δ18O records from Kuhai Lake[24], f: Autogenic carbonate δ18O records from Lake Bande[25], g: Silt content records from Nir’pa Co sediments[26], h: Temperature reconstruction sequences from Qinghai Lake[47], i: Temperature reconstruction sequences from Tiancai Lake[27].
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图 5 Ⅲ区过去10 ka青藏高原古气候记录对比
a: 纳木错Kclr记录[48]，b: 纳木错Ca记录[49]，c: 纳木错Ti记录[49]，d: 色林错Ti记录[51]，e: 色林错自生碳酸盐δ18O记录[52]，f: 兹格塘错植被类型重建记录[29]。
Figure 5. Comparison of paleoclimate records in Zone III from the past 10 ka in the Tibetan Plateau
a: Nam Co Kclr records[48], b: Nam Co Ca records[49], c: Nam Co Ti records[49], d: Seling Co Ti records[51], e: Autogenic carbonate δ18O records from Seling Co[52], f: Vegetation type reconstruction records from Zigetang Co[29].
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图 8 类NAO、类SAM以及类PDO条件百年尺度的指数变化
a: 5 kaBP以来的类NAO条件指数[105]，b: 6 kaBP以来的类SAM条件春季（3—5月，MAM）指数[90]， c: 10 kaBP以来的类PDO条件指数[106]。
Figure 8. Changes in centennial-scale indices of the NAO-like, SAM-like, and PDO-like scenarios
a: NAO-like index since 5 kaBP[105], b: SAM-like index for spring (March to May, MAM) since 6 kaBP[90], c: PDO-like index since 10 kaBP[106].
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图 7 NAO、IOD以及ENSO对青藏高原的湿度影响示意图
Figure 7. Schematic diagram of the impacts of NAO, IOD, and ENSO on the humidity of the Tibetan Plateau
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图 6 11 kaBP以来大尺度海气耦合系统指数与北半球古降水/水分变化以及其他记录的比较
a: 加拉帕戈斯群岛El Junco湖的含砂量记录，可视为赤道太平洋东部厄尔尼诺活动的代表[69]；b: 智利南部Lago CipresesLC: 湖泊记录中Pilgerodendron uviferum丰度的百分比，可解释为类SAM条件负相位条件的代表[84]；c: 火山影响活动[85]；d: Cariaco盆地沉积物中的Ti记录[78]；e: 北半球中高纬度温度变化重建记录[83]，f: 亚洲中部干旱区湿润指数[54]；g: 冰川活动记录[54]，h: 纳木错Kclr记录[48]；i: 两条董哥洞δ18O记录[80-81]。
Figure 6. Comparison between large-scale ocean-atmosphere coupling system indices and the Northern Hemisphere paleoprecipitation / moisture changes and other records since 11 kaBP
a: Sediment sand content records from El Junco Lake in the Galápagos Islands, representing eastern equatorial Pacific El Niño activity[69, b: Percentage of Pilgerodendron uviferum abundance in lake records from, interpreted as a representation of SAM-like negative phase conditions[84]; c: Volcanic influence activity[85]; d: Ti records in sediments from the Cariaco Basin[78], e: Reconstruction records of temperature changes in the mid to high latitudes of the Northern Hemisphere[83]; f: Humidity index for the arid regions of central Asia[54]; g: Glacier activity records[54]; h: Nam Co Kclr records[48];i: Two carbonate δ18O records from Dongge Cave[80-81].
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