Preliminary study on Oligo-Miocene hydrological changes in Southeast Asia and their driving mechanisms
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摘要:
新生代印尼海道的启闭对印度-太平洋暖池演化和大气环流模式变迁有重大影响。然而,受限于构造和古环境重建资料的缺乏,这三者之间的逻辑关系和驱动机制还缺乏清晰的图景。本文梳理了孢粉记录、煤层沉积、浅海碳酸盐沉积和生物地理演化等方面的证据,提出东南亚水文气候在渐新世与中新世之交发生重大调整的认识,即从渐新世的相对干旱条件转型为贯穿整个中新世的持续湿润状态。结合最近的模拟研究,认为东南亚水文气候演变同时受到全球因素和区域构造要素的影响。渐新世与中新世之交和中中新世晚期至晚中新世早期,印尼海道的持续关闭可以通过限制太平洋-印度洋次表层水的交换,进而扩大太平洋一侧的温跃层深度以及经纬向的海表温度梯度,进一步增强沃克环流,最终可能促使东南亚在渐新世与中新世之交发生了干湿格局的转换,并抵消了中中新世晚期至晚中新世全球降温对区域水文气候的影响。目前的研究仍存在不确定性,未来亟需更多的地质记录和模拟研究来准确厘定海道关闭-暖池演化-大气环流之间的联系。
Abstract:The closure of the Indonesian Seaway played a key role in the evolution of the Indo-Pacific Warm Pool and associated atmospheric circulation during the Cenozoic. However, the relationship between the closure of the seaway, the evolution of the warm pool, and the shift in atmospheric circulation remains unclear due to poor constraints in tectonic and paleoenvironmental reconstructions. This study reviews the historical literature, including evidence from pollen records, coal deposits, shallow marine carbonate deposits, and biogeographic evolution. The results show that the hydroclimate in Southeast Asia underwent significant changes during the Oligo-Miocene transition, shifting from relatively dry conditions in the Oligocene to persistently wet conditions throughout the Miocene. Combined with recent simulation studies, it was concluded that the hydrological changes in Southeast Asia were influenced by both global and regional factors. The narrowing and closure of the seaway may have increased the gradient between the east-west thermocline depth and the east-west sea surface temperature in the Pacific Ocean, limiting the exchange of subsurface water between the Pacific and Indian Oceans. This in turn led to a strengthening of the Walker Circulation, which subsequently induced hydrological changes in Southeast Asia after the Oligo-Miocene boundary and mitigated the effects of global cooling over the Late Miocene. Uncertainties remain in current studies, and more geological records and simulation studies in the future would help to accurately characterize the relationship between seaway closure, warm pool evolution, and atmospheric circulation in the Oligo–Miocene.
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印度-太平洋暖池(简称印太暖池)作为地球的“热引擎(heat engine)”和“蒸汽引擎(steam engine)”,其海表温度(SST)常年大于28℃并伴随有强烈的大气深对流活动(图1a)[1]。根据克劳修斯-克拉伯龙方程,饱和水蒸气压与SST呈指数关系,导致大气深对流活动强度对暖池SST响应十分敏感,暖池SST的微弱调整能通过与大尺度环流(即沃克环流和哈德雷环流)相互作用影响整个气候系统,尤其是水文循环[2]。印尼海道是目前唯一的低纬开放大洋通道,其在新生代的逐步收缩和关闭是暖池地质演化的重要制约因素。一方面在地理空间上为近地表东风驱动下的暖水堆积提供物理屏障;另一方面通过调控印尼贯穿流进而影响暖池的发育并最终作用于全球水文循环[3]。
图 1 印太暖池现代海流与本文研究所引用的站位资料a:印太暖池分布范围及其表层洋流,其中橙色实线代表暖流,蓝色实线代表寒流,绿色实线代表季风淡水输入(修改自文献[37–38]);b:东南亚代表性钻井孢粉剖面(黑色圆点, 修改自文献[39–43])和主要含煤盆地分布(陆上褐色区域, 修改自文献[44–45]),华莱士生物区及其界限(白色虚线, 修改自文献[46])。Figure 1. Modern oceanography in the Indo-Pacific Warm Pool and sites cited in this studya: Distribution of the Indo-Pacific Warm Pool (black lines) and surface oceanic circulation in the region. Warm currents (orange lines), cold currents (blue lines), and monsoonal freshwater plumes (green lines) are presented (modified after references [37–38]); b: location of palynological profiles from representative wells (black dots, modified after references [39–43]) and main coal-bearing basins of Southeast Asia (brown patches on land, modified after references [44–45]), the Wallace’s zoogeographical region and its boundaries are denoted (white dashed lines, modified after reference [46]).印尼海道收缩关闭导致印尼贯穿流的供给水团由暖-咸的南太平洋水转变成相对冷-淡的北太平洋水,前者折返进入北太平洋赤道逆流促进西太平洋表层—次表层水体的暖化,后者随印尼贯穿流进入印度洋南赤道暖流导致印度洋表层—次表层水体的冷却(图1a)[4–5]。同时,更多的亚南极模态水和南极中层水通过赤道潜流在东赤道太平洋上升,导致温跃层浅化和冷水上涌,强化了太平洋沃克环流,并进一步促进暖池的发育[6–8]。印尼海道关闭时间一直存在激烈争议,部分学者认为在渐新世与中新世之交(Oligocene–Miocene transition, OMT; 约23 Ma),澳大利亚板块与东南亚陆块碰撞,印尼海道便开始收缩关闭[9–11]。此时,太平洋与印度洋之间深部水体交换可能已经受到显著限制,导致太平洋底栖有孔虫稳定氧(δ18O)、碳(δ13C)同位素明显低于印度洋[10,12]。然而,印尼海道的收缩关闭与暖池形成之间的对应关系存在较大的争议。例如,Von Der Heydt和Dijkstra[13]基于模拟指出,印尼海道的相对开放并不会阻止暖池的形成,只是海道关闭前暖池的热中心可能位于印度洋。
其他学者认为印尼海道关键性闭合时段分别在中中新世晚期—晚中新世早期(约12~10 Ma)[14–18]或者上新世晚期(约4~3 Ma)[5,19–20]。这些闭合过程逐步推动了暖池向现代格局的过渡和发育。印尼海道关闭过程中,由于暖水的积聚导致暖池温跃层增厚,混合层浮游有孔虫丰度相对温跃层种属增加,垂向上浮游有孔虫氧同位素梯度减小,并出现经向温跃层深度梯度[15–16]。太平洋经向与纬向温跃层深度梯度在上新世同时显著增加,也意味着在印尼海道的最后关闭阶段,现代暖池格局正式形成[14–15,21]。
暖池的形成与强化会显著重塑广大区域的水文循环格局。首先,暖池强化与太平洋沃克环流增强密切相关。太平洋沃克环流上升支与下沉支分别对应着东南亚西部和南美地区,两个地区的降雨分别出现增加和减少[22–23]。通过行星波传播,暖池强化也会增强印度洋沃克环流,进一步促进东非干旱[5]和南亚次大陆湿润化[24]。太平洋沃克环流的强化反过来也可以维持东西太平洋的热不对称性,同时促进哈德雷环流的增强[8,24–25],并导致热带辐合带和西太平洋副热带高压分别向南撤和向西延伸,这些会强化亚洲内陆地区的干旱化[26–27]。此外,澳大利亚西部地区也会因暖池强化出现干旱的趋势[20,28–29]。
由于暖池发育和区域水文循环演变之间存在紧密联系,本文通过综述东南亚地区渐新世至中新世相关古气候重建资料,探讨东南亚干湿气候格局的时空演变与印尼海道闭合以及暖池发育的耦合关系,并结合最近的模拟研究,尝试探讨三者之间的物理驱动机制。
1. 地质背景
1.1 全球气候变化
新生代全球气候逐渐变冷的进程中,地球系统自渐新世从两极无冰的暖室期进入南极单极冰盖发育的冰室期。渐新世期间,深海底栖有孔虫δ18O为1‰~3‰,全球均温高于现在约3~4℃,南极冰盖大幅扩张,海平面高约30~50 m(图2a–c)[30–32]。大气二氧化碳(CO2)浓度在(300~700 )× 10−6之间波动(图2d)[31]。SST变化呈显著空间异质性,高纬海域SST为15~20℃,中低纬海域SST约为30℃(图2e)[33]。渐新世气候演化根据温度变化大致可以分为3个阶段:始新世—渐新世之交后,全球经历了约2.5 Ma的温度回升,随后发生约5 Ma的长期降温,之后再次出现约2.5 Ma的回暖。
图 2 全球气候变化和印尼海道构造演变a:全球深海底栖有孔虫δ18O[30];b:全球平均地表温度相对变化[31];c:全球海平面相对变化[32];d:大气CO2浓度[31];e:基于TEX86重建的全球SST,红色代表中低纬度地区,蓝色代表高纬地区[33];f:印尼海道构造演变重建(修改自文献[34,38])。Figure 2. Global climate change and tectonic evolution of the Indonesian Seawaya: Global deep-sea benthic foraminiferal δ18O[30]; b: variation in global mean surface temperature relative to the present estimated from benthic δ18O[31]; c: variation of sea level relative to the present estimated from benthic δ18O[32]; d: atmospheric CO2 estimate[31]; e: global SST estimated from TEX86, where red line represents middle-low latitude region and blue line represents high latitude region[33]; f: tectonic reconstruction of the Indonesian Seaway (modified after references [34,38]).OMT期间,深海底栖有孔虫δ18O增加约1‰,全球均温下降约0.5℃,南极冰盖再次快速扩大,海平面下降约10 m。同时,大气CO2浓度下降约100×10−6。高纬海域SST下降约5℃,中低纬海域SST变化仍不显著。OMT之后,整个中新世期间,气候波动较渐新世显著。深海底栖有孔虫δ18O为1‰~4‰,全球均温比现在高约1~5℃,海平面高约10~60 m。大气CO2浓度为(300~500) × 10−6,与现代较为接近。高纬海域SST为10~20℃,中低纬海域依然在30℃左右。中新世气候演化大致可以分为两个阶段:OMT之后,全球经历了约8 Ma的温度回升,随后进入长达约10 Ma的温度下降。
1.2 印尼海道构造演变
印尼海道的收缩关闭过程受板块运动的显著影响[9,34–36,38]。中始新世(约45 Ma)起,澳大利亚板块开始北移,向东南亚陆块俯冲。这一运动引发了东南亚陆块内部的一系列构造变形。其中,东南部的断裂活动形成了望加锡海峡,将婆罗洲与西苏拉威西分离。该海峡于渐新世逐渐成为深水通道,为北太平洋中层水和表层水进入印度洋提供通路。这一阶段,澳大利亚板块尚未与东南亚陆块接触,南太平洋水也能自由进入印度洋(图2f)。
澳大利亚板块继续北上,直至OMT,苏拉威西岛支(Sula Spur)与北苏拉威西弧碰撞,导致婆罗洲、苏拉威西和班达岛弧发生造山运动。同时,新几内亚被动边缘与菲律宾-哈马黑拉-南卡罗林岛弧碰撞。一系列碰撞事件导致澳大利亚板块与东南亚陆块之间的深水通道关闭,形成了华莱士区,并促进了大量浅海碳酸盐发育。印尼海道开始收缩关闭,太平洋与印度洋深水水体交换因此开始受阻。
随着澳大利亚板块的继续俯冲,婆罗洲等微型陆块发生逆时针旋转。中中新世(约15 Ma)以来,爪哇俯冲带回滚后撤至班达湾(Banda embayment),导致苏拉威西拉伸活动广泛发育,加速了陆地隆起和面积的增加,进一步促进印尼海道收缩。直到早上新世(约5 Ma),印尼群岛的进一步出露,尤其是哈马黑拉的快速隆升,显著地减小了印尼海道规模,印尼海道随之完成闭合。
2. 干湿气候格局的时空演变
本文从孢粉记录、煤层沉积、碳酸盐沉积和生物地理演化4个研究程度较高的角度,尝试对东南亚渐新世至中新世干湿气候格局时空演变进行回顾和总结。表1为本文主要引用资料的基本信息。
表 1 主要引用资料基本信息Table 1. Basic information of main references cited for this study站位名称 地点 替代性指标 气候指示意义 覆盖时间 地层方法 参考文献 Well C 东南亚越南南部 孢粉组合相对丰度 植被类型干湿变化 24.2~28.2 Ma 生物地层 [42] Well E 东南亚越南南部 孢粉组合相对丰度 植被类型干湿变化 20.9~23.4 Ma 生物地层 [42] Belut-3 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 33.7~34.8 Ma 生物地层 [39] Kambing-1 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 27.5~29.4 Ma 生物地层 [39] Kambing-1? 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 23.8~27.5 Ma 生物地层 [39] Kadal-1 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 21.2~23.8 Ma 生物地层 [39] Bergading Deep-3 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 7.7~16.9 Ma 生物地层 [43] Dengkis-1 东南亚婆罗洲北部 孢粉组合相对丰度 植被类型干湿变化 6.6~14.2 Ma 生物地层 [43] Kuda Laut-1 东南亚婆罗洲北部 孢粉组合相对丰度 植被类型干湿变化 3.5~23.1 Ma 生物地层 [43] Bukoh-1 东南亚婆罗洲北部 孢粉组合相对丰度 植被类型干湿变化 20.2~28.2 Ma 生物地层 [43] Well A 东南亚婆罗洲东部 孢粉组合相对丰度 植被类型干湿变化 0~8.5 Ma 生物地层 [40] Well B 东南亚婆罗洲东部 孢粉组合相对丰度 植被类型干湿变化 0~9.4 Ma 生物地层 [40] Well X 东南亚爪哇海东部 孢粉组合相对丰度 植被类型干湿变化 23~30 Ma 生物地层 [41] Well Y 东南亚爪哇海东部 孢粉组合相对丰度 植被类型干湿变化 23~30 Ma 生物地层 [41] 东南亚 煤层分布 降雨强度干湿变化 新生代 [44–45] 东南亚 浅海碳酸盐生物相相对面积 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐生物相面积相对占比 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐台地数量 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐浅滩/建隆数量 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐格架礁分布 降雨强度干湿变化 新生代 [47] 东南亚 降水量数值模拟 降雨强度干湿变化 0~30 Ma [46] 巴基斯坦 食草动物牙釉质碳同位素 植被类型干湿变化 0~33 Ma [48] 巴基斯坦 食草动物牙釉质氧同位素 降雨强度干湿变化 0~33 Ma [48] 中国西北 孢粉组合相对丰度 植被类型干湿变化 0~45 Ma [49] 中国西北 孢粉组合相对丰度 植被类型干湿变化 0~39 Ma [50] 非洲东西沿海 有机质叶蜡烷烃碳同位素 植被类型干湿变化 0~23 Ma [51] 非洲埃塞俄比亚 有机质叶蜡烷烃碳同位素 植被类型干湿变化 渐新世-早中新世 [52] IODP U1464 澳大利亚西北 沉积物钾元素含量 降雨强度干湿变化 5~15 Ma 生物地层 [53] 全球 底栖有孔虫氧同位素 全球冰量底层海水温度 新生代 [30] 全球 底栖有孔虫氧同位素 全球平均温度 新生代 [31] 全球 多指标(例如浮游植物长链不饱和酮类
碳同位素、硼同位素等)大气CO2浓度 新生代 [31] 全球 底栖有孔虫氧同位素 全球海平面 0~40 Ma [32] 全球 有机质TEX86 SST 新生代 [33] 2.1 孢粉组合
孢粉的外壳具有较强的抗风化作用而容易保存,并且体积小、质量轻、产量大、分布范围广,因此孢粉可以有效重建地质历史时期植被分布和气候环境演变。Robert J. Morley等学者通常将东南亚孢粉序列划分为雨林、荒原森林、泥炭沼泽、河岸带、海岸带、季节性干旱和山地组合,其中部分组合对气候干湿具有指示意义。例如,雨林组合(Dipterocarpaceae, Calamus)和泥炭沼泽组合(Alangium, Blumeodendron, Campnosperma, Ilex)反映湿润的气候;荒原森林组合(Casuarina, Dacrydium)反映极端湿润的气候;季节性组合(Poaceae, Pinus, Compositae)则反映干旱的气候。
东南亚在渐新世整体呈现干旱气候,主要以季节性组合为主。到了晚渐新世与早中新世之交,气候由干旱转向湿润,雨林和泥炭沼泽组合取代季节性组合,并出现荒原森林组合。早中新世至晚中新世期间,湿润气候条件得以维持,雨林和泥炭沼泽组合继续占据主导地位,部分地区在晚中新世出现少量季节性组合。以代表性钻井为例(图1b、图3),越南南部九龙盆地钻井(Well C和Well E)[42]显示,晚渐新世早期至晚期,季节性植被逐渐增加,到早中新世早期减少,并逐渐被雨林植被取代。马来半岛附近的马来盆地钻井(Bergading Deep-3)[43]显示,早中新世晚期到晚中新世,季节性植被占比极小,整体以雨林和泥炭沼泽植被为主。西纳土纳盆地钻井(Kambing-1, Kadal-1, Belut-3)[39]显示,早渐新世以季节性植被为主,晚渐新世季节性植被持续出现,雨林和泥炭沼泽植被逐渐增加;晚渐新世末至早中新世,季节性植被逐渐消失,雨林和泥炭沼泽植被逐渐占主导地位,同时伴随荒原森林植被的出现。婆罗洲西北部曾母盆地钻井(Bukoh-1, Kuda Laut-1)相较其他站位,自中新世以来都有明显的季节性组合出现,这可能受到北部地区季节性孢粉输入的影响[43]。
文莱-沙巴盆地钻井资料(Dengkis-1)[43]与马来盆地相似,中中新世至晚中新世,没有明显的季节性组合,以雨林和泥炭沼泽植被为主,并伴随荒原森林植被的频繁出现。婆罗洲东部Tarakan盆地(Well A)和Kutai盆地钻井(Well B)[40]也显示,晚中新世以来几乎无季节性植被,以雨林和泥炭沼泽植被为主,可持续到早更新世。婆罗洲北部陆上沙捞越盆地早中新世Balingian组煤层发育大量以木麻黄属(Casuarina)和陆均松属(Dacrydium)为代表的荒原森林植被类型[54],早上新世Begrih组煤层发育大量的泥炭沼泽植被类型[55],也说明婆罗洲潮湿的气候条件从早中新世持续到早上新世。此外,沙捞越盆地的Nyalau组地层中,晚渐新世地层频繁出现禾本科植被,到早中新世则出现大量的木麻黄属和泥炭沼泽植被,说明晚渐新世至早中新世气候由干旱向湿润转型[56]。
不同的是,爪哇和苏门答腊群岛干湿转换的时间发生于早渐新世与晚渐新世之交。爪哇海东部钻井资料(Well X和Well Y)[40]显示,晚渐新世发育大量的泥炭沼泽和荒原森林植被,几乎没有季节性植被。爪哇海西部的钻井[57]中,早渐新世地层季节性植被(禾本科和莎草科)出现频繁,晚渐新世地层泥炭沼泽植被显著增加。相似地,苏门答腊中部的陆上钻井也显示,早渐新世禾本科植被经常出现,晚渐新世泥炭沼泽植被大量发育的情况[57]。
2.2 煤层沉积
煤层沉积的形成同时需要构造条件和适宜的气候环境。煤层通常出现在温暖湿润的环境中,自中生代以来,约50%的煤层发育在年降水为700~
1200 mm的区域[58]。东南亚地区,苏门答腊与爪哇地区的煤层主要形成于晚渐新世至早中新世早期,而婆罗洲及其东侧的煤层主要形成于中中新世(1b、图4)。这些时期的湿润气候条件与孢粉组合记录一致。需要注意的是,并非所有干旱时期都伴随煤层的缺失。煤的灰分含量和硫分含量通常能够揭示泥炭沼泽的类型。高位泥炭沼泽,中央部位地势比边缘稍高,主要依赖降雨补给,营养贫乏,偏酸性,泥炭藓为优势植被,灰分含量和硫分含量低;低位泥炭沼泽,中央部位地势比边缘部位稍低或持平,由于存在地表水和地下水补给,营养丰富,偏弱酸性—弱碱性,草本植物为优势植被,灰分含量和硫分含量高。低位泥炭沼泽通常被认为是高位泥炭沼泽的初期形态,当降雨丰沛且长期稳定的情况下,低位泥炭沼泽会向高位泥炭沼泽转化[59–60]。因此,相较于低位泥炭沼泽,高位泥炭沼泽倾向出现在季节性弱的湿润气候条件下。渐新世至中新世东南亚地区煤的灰分含量和硫分含量普遍较低,中值分别为2.24%和0.19%(图5a)。这与德国Lower Rhine盆地(灰分含量<10%, 硫分含量<1%)[61]和哥伦比亚Cesar-Rancheria盆地(灰分含量<6%, 硫分含量<1.1%)[62]等地的高位泥炭沼泽相似。
煤的显微组分也可以用于指示成煤环境。镜质体一般是在湿润/还原条件下,由植物遗体在凝胶化作用过程中转化形成,而惰质体则是野火或氧化作用的产物[76-77]。禾本科植物的扩张和季节性干旱的增强可能导致高降雨地区的野火频率上升[78],并且干旱低水位条件下氧化作用的增强也会加强惰质体的生成。凝胶化指数(gelification index, GI)即凝胶化组分与非凝胶化组分之比,反映了泥炭积累过程中泥炭沼泽的湿润程度和持续时间,通常湿润环境下GI值较高。由于GI值计算使用极不统一,为方便评估,本文根据镜质体与惰质体之比重新计算GI值[79]。当GI值大于1时,表明成煤环境处于长期湿润的条件下。渐新世至中新世东南亚煤层显示出高镜质体含量(中值为85.5, 图5b),极低的惰质体含量(中值为3, 图5b),以及较大的GI值(中值为26.1, 图5c),这些特征与灰分含量和硫分含量的数据一同表明,该地区的煤层均形成于湿润的气候背景之下。
2.3 碳酸盐沉积
碳酸盐沉积物的生产对环境条件非常敏感,降雨引起的淡水和陆源碎屑及营养盐输入会影响水体的盐度、水质和营养盐含量,从而影响碳酸盐生产者的类型[80–82]。不同碳酸盐生产者对盐度的耐受范围不同,例如造礁珊瑚耐盐范围为25‰~42‰,底栖有孔虫耐盐范围为30‰~45‰。虽然盐度的变化本身不足以影响碳酸盐沉积,但它与CO2以及温度的共同作用会导致文石的饱和度降低。在近岸区域,陆源物质输入会增加水体中的悬浮物浓度,导致透光层变窄。此外,营养盐的输入还可以导致本来贫营养的热带海水富营养化,造成藻类勃发,进而使上层水体透明度下降。这些因素都会抑制碳酸盐的形成,并导致寡光型碳酸盐生产者大量发育,使碳酸盐沉积无法形成建造格架。
新生代超过2/3的浅海碳酸盐沉积发生在东南亚。在渐新世时期,较为强烈的季节性气候条件使得淡水和陆源物质的输入频繁变动。相对于稳定的湿润气候,这些条件对碳酸盐的沉积不利。由于底栖有孔虫对陆源物质输入具有更强的耐受性,而且它们的寿命较短,能更好地适应季节性光照不足的环境,因此渐新世出现较为广泛的寡光型碳酸盐堆积[47,83]。OMT之后,东南亚气候由干旱转向湿润,导致碳酸盐沉积生物相从以底栖有孔虫为主转向以文石珊瑚为主。这一转变伴随碳酸盐沉积面积的显著增加,碳酸盐台地、碳酸盐浅滩和建隆出现不同程度的扩张,并且格架礁广泛发育(图6a–f)。尽管在早中新世之后碳酸盐沉积面积减少,但珊瑚主导的碳酸盐沉积比例直至晚中新世并未发生显著改变,这表明东南亚地区在整个中新世期间保持了相对稳定的湿润状态。
图 6 渐新世至中新世东南亚水文气候变化与浅海碳酸盐沉积发育情况对比a:基于孢粉组合重建的水文气候变化,b:赤道不同碳酸盐生物相的相对面积,c:赤道不同碳酸盐生物相相对面积的相对比例,d:碳酸盐台地数量,e:碳酸盐浅滩和建隆数量,f:格架礁的发育情况(修改自文献[47,84])。Figure 6. Comparison of hydrological changes in the Oligocene to Miocene with the development of shallow-marine carbonate formations in Southeast Asiaa: Reconstructed hydrological changes based on palynological assemblages, b: the relative areas of equatorial carbonate of different carbonate biofacies, c: the relative proportion of the relative areas of equatorial carbonate to different carbonate biofacies, d: numbers of carbonate platforms, e: numbers of carbonate shoals/buildups, f: development of framework reefs(modified after Refs. [47,84]).2.4 生物地理演化
东南亚是现今地球的生物多样性热点地区。始新世时期,热带雨林分布扩展至中纬度地区。渐新世以来,随着全球降温和干旱化,热带雨林大幅度向低纬度地区退缩。由于东亚大陆与赤道地区相连,东南亚成为许多北方热带植物区系的重要避难所[85]。从早中新世至中中新世早期,温暖湿润的气候条件促使热带雨林再次向亚热带扩张,随后在中中新世晚期至晚中新世,雨林再次退回到低纬度地区[85–86]。龙脑香科(Dipterocarpaceae)是东南亚热带雨林的主要植物之一,起源于非洲大陆,随后扩散至东南亚。龙脑香科在约20.4 Ma后属种分化速率显著增加,例如娑罗属(Shorea)的多样性在婆罗洲等地显著增加,指示早中新世之后东南亚出现湿润的气候格局[86]。
生物源区季节性干旱气候的形成可能导致生物多样性减少,并降低生物的迁徙频率。Klaus等[87]发现,东南亚生物向印度次大陆的迁徙频率在24~21 Ma期间出现停滞,之后逐渐增加并在中中新世时期达到峰值,在晚中新世时又出现下降。以上可能反映了东南亚早中新世至中中新世湿润气候的扩张,而在晚中新世期间,东南亚北界季节性干旱增强,导致东南亚与印度次大陆之间生物交换频率降低。然而,东南亚和澳大利亚之间的生物交换与前者不同。穿越华莱士生物区是两地生物交换的必经之路(图1b)。Skeels等[46]通过模型重建指出,类似华莱士生物区的湿润气候在东南亚出现的时间约为20或30 Ma,与热带雨林扩张的时间一致。持续的湿润气候导致东南亚生物在华莱士生物区定殖频率持续增加,直到上新世达到峰值[46,88]。不同的趋势可能反映了东南亚气候变化存在空间异质性。
两栖动物中的树蛙对海水不耐受,扩散能力较弱,其种群数量跟森林面积和降雨强度存在紧密联系,因此是反映东南亚干湿气候格局演化的有利工具[89–91]。研究发现,树蛙科泛树蛙属(Rhacophoridae, Polypedates)起源于晚渐新世的巽他群岛,中中新世时它的分布范围向北扩散至东南亚大陆和华南地区,晚中新世至上新世从东南亚大陆又收缩回巽他群岛[91]。这一发现进一步表明,东南亚气候变化存在显著的空间异质性,推测晚中新世以来其北界变得相对干燥,而中心主体地区则维持湿润状态。
综合以上资料,东南亚绝大部分地区在渐新世时期表现为相对干旱的气候条件,而在OMT发生了气候转型,由干旱逐渐转变为湿润,并持续整个中新世(图7a)。这种干湿格局的转变得到了部分模拟研究的支持。例如,Skeels等[46]利用HadCM3L耦合模式模拟了近30 Ma的降水情况,显示自OMT以来的整个中新世,东南亚降水在持续增加(图7a)。类似地,Zhu等[92]利用CESM耦合模式重建了新生代以来东南亚的降水历史,结果显示晚渐新世以来降水呈阶梯式增长,早中新世早期至晚中新世早期降水量相对稳定,晚中新世早期之后降水继续增加。尽管这些模拟结果存在些许差异,但基本上表明东南亚在渐新世相对干旱,在渐新世与中新世之交经历了明显的干湿格局转型,并在中新世维持相对湿润的气候条件。
图 7 渐新世至中新世东南亚与其他区域的水文气候对比a:基于孢粉组合[84]和模拟[46]重建的东南亚水文气候变化;b:印尼海道演变历史[38,34];c:全球深海底栖有孔虫δ18O[30];d:全球平均地表温度相对变化[31]和海平面相对变化[32];e:大气CO2浓度[31];f:基于TEX86重建的全球SST,红色为中低纬地区,蓝色为高纬地区[33];g:巴基斯坦大型食草哺乳动物牙釉质δ18O和δ13C[48];h:亚洲内陆旱生植物占比,其中绿色散点为Ephedra + Nitraria + Chenopodiaceae + Artemisia之和[49,108],橙色柱状为干草甸和沙漠景观植被占比之和[50];i:非洲叶蜡烷烃δ13C[51–52];j:全球晚中新世C4植被扩张开始时间[51];k:澳大利亚西北部IODP U1464钾元素含量,大于0.3为湿润,小于0.2为干旱[53]。Figure 7. Comparison of hydrological changes in the Oligocene to Miocene in Southeast Asia with other regionsa: Reconstructed hydrological changes based on the palynological assemblages[84]and model stimulation[46]; b: the evolution history of the Indonesian Seaway[38,34]; c: global deep-sea benthic foraminiferal δ18O[30]; d: global mean surface temperature[31] and sea level change[32] relative to the present estimated from benthic δ18O; e: atmospheric CO2 estimate[31]; f: global SST estimated from TEX86, where red line represents middle-low latitude region and blue line represents high latitude region[33]; g: δ18O and δ13C of tooth enamel bioapatite for large herbivorous mammal in Pakistan[48]; h: xerophytic vegetation changes in Central Asia, where green dots represent the sum of Ephedra + Nitraria + Chenopodiaceae + Artemisia[49,108], and orange bars represent the sum of steppe and desert vegetation[50]; i: compilation of δ13C for plant wax n-alkanes in Africa[51–52]; j: the onset of C4 vegetation expansion in the Late Miocene across the globe[51]; k: potassium content of the IODP U1464 located in the northwestern Australia, with wet and dry conditions being represented by >0.3 and <0.2, respectively[53].3. 全球和区域对比
渐新世时期,印尼海道尚未发生收缩关闭(图2f、图 7b),东南亚干湿变化与深海底栖有孔虫δ18O、全球平均温度以及海平面变化基本同步。这种同步表现为:δ18O降低时,温度上升,海平面升高,东南亚气候变湿;相反,δ18O增大时,温度下降,海平面降低,东南亚气候变干(图7a、 c–d)。然而,东南亚干湿变化与大气CO2浓度以及SST变化未呈现对应的变化关系(图7a、 c、 e–f)。
OMT以来,印尼海道开始收缩关闭,早中新世时期,底栖有孔虫δ18O、全球平均温度和海平面情况与渐新世的暖期相似。但是,此时东南亚已经呈现相对较强的湿润状况。中中新世,各气候环境参数达到峰值,全球进入大暖期,东南亚继续保持湿润的状况。中中新世晚期至晚中新世,底栖有孔虫δ18O持续增加,温度持续降低,海平面持续下降。相应地,大气CO2浓度持续下降,高纬SST持续下降。尽管如此,东南亚地区仍然维持着湿润的状态。
OMT时期的干湿气候格局转型并不是东南亚独有的区域现象,不同地区都有类似的气候变迁现象,但表现不一。与东南亚一致的是,巴基斯坦经历了由干旱向湿润环境的转变,大型食草哺乳动物牙釉质δ18O开始减小(图7g)[48]。但OMT之后至晚中新世,亚洲内陆经历了显著的干旱化[93–94]。在中国西北地区,陆地景观开放程度增加,森林植被由干草原和沙漠植被取代,湿润环境永久转化为干旱-半干旱环境(图7h)[49–50]。由于亚洲内陆的干旱化,亚洲内陆的风尘向西北太平洋的输入通量显著增加[95–96]。相似的情况出现在非洲,叶蜡烷烃δ13C由下降转为上升趋势,表明气候逐渐干旱并导致C4植被在陆地生态系统中的占比逐渐增加(图7i)[51–52]。
晚中新世的全球大范围干旱化现象尤为显著,最典型的标志就是C4植被的显著扩张(图7j)。C4植被的显著扩张最早发生于北非和东非(约10 Ma)[51],随后出现在南非(约8 Ma)[97],巴基斯坦、印度和尼泊尔等南亚地区(约9.5~7.4 Ma)[48,98],中亚和东亚(约8~5.3 Ma)[50,99–100],南美(约7 Ma)[101]和北美(约7 Ma)[102–103]。与上述地区的干旱化趋势不同,东南亚整体未出现显著的干旱化,气候状态保持较为稳定的湿润条件。有趣的是,澳大利亚西北部也保持了较为稳定但整体干旱的气候状态(图7k)[29,53,104]。通过与全球气候以及不同区域水文气候演变记录的对比,可以认为渐新世至中新世东南亚干湿气候格局的转变不仅受全球信号的影响,而且也受区域因素的调控,因此展现出区域独特性。
4. 东南亚气候变迁的驱动机制
渐新世时期,东南亚干湿气候变化与全球温度和海平面变化基本同步,说明该时期水文循环可能与温度-冰盖介导的水热反馈有关。OMT之后,东南亚从渐新世的相对干旱转向湿润。该水文气候格局的转型与澳大利亚板块-东南亚陆块开始碰撞的时间相近,因此二者之间可能存在动力学机制上的联系[10,84]。如前所述,碰撞导致印尼海道收缩关闭以及印太暖池现代格局的逐步成型,可能最终影响了东南亚地区的大气环流模式。
Zhang等[105]利用NorESM-L模式,在Scotese[106]的地形工作基础上,重建了新生代印太暖池的变化。研究显示,印太暖池中心从早始新世至晚中新世始终在印度洋一侧(约60°~90°E),直至晚上新世才向东移动至太平洋印尼海域。这与Von Der Heydt和Dijkstra[13]基于CCSM模式重建的早中新世暖池位置相似。同时,印太暖池的面积和强度从早始新世至晚中新世持续减小和下降。以上可能说明,OMT时期印尼海道的收缩关闭并未通过调节暖池的位置和强度变化影响大气环流模式,并导致东南亚水文气候格局转型。这一点得到一项关于太平洋沃克环流模拟重建工作的支持[107]。太平洋沃克环流的上升支从早始新世至晚中新世维持在约90°~120°E之间,且强度整体呈上升趋势。
然而,海道的收缩关闭可能通过限制次表层水体影响大气环流模式。根据Zhang等[105]的模拟输出结果,水深150 m处的海水温度在早中新世发生了重要的格局转变。早渐新世时印度洋和太平洋两侧次表层水体交换通畅,水温分布一致;而早中新世时印度洋次表层水体显著降温,海温分布已经与工业革命前情景相似。这说明海道关闭并非从上新世才产生影响,海道收缩关闭早期主要限制较深处的水体交换,导致西太平洋温跃层厚度大于东印度洋[10]。Karas等[6]曾指出上新世印尼海道关闭并未导致东印度洋SST发生显著的变化,反而是东印度洋次表层水温度下降了4℃。这一情景与Zhang等[105]的模拟结果非常相似。印尼海道的关闭导致了更多冷-淡的北太平洋次表层水进入印度洋,促使印度洋温跃层发生降温和浅化(图8)。同时,更多暖-咸的南太平洋次表层水进入赤道逆流,并且南大洋水团通过赤道潜流在东太平洋上升,这些导致太平洋沃克环流的增强[6–8]。该猜想得到最近东西太平洋温跃层深度重建工作的支持,Matsui等[109]发现早中新世期间(18.8~16.6 Ma)东西太平洋温跃层倾斜程度整体较陡。因此,OMT时期的海道收缩关闭,可能通过限制次表层水体的变化,强化太平洋沃克环流,促进东南亚由干转湿。
中中新世早期(16.5~13.8 Ma),大气CO2浓度增加至中新世峰值,全球温度也发生明显升温,太平洋东西向温跃层倾斜程度相较于早中新世变缓[109],说明中中新世暖期的湿润气候条件不是沃克环流强化的结果,更可能与该时段高大气CO2浓度导致的大气持水能力增加有关。中中新世晚期至晚中新世,大气CO2浓度与全球温度均显著下降,东南亚并未因此出现与其他地区类似的干旱事件。中中新世晚期(13.07~11.9 Ma),太平洋东西向温跃层倾斜程度恢复至早中新世期间的状态[109]。晚中新世两大洋次表层海水温度差异与早中新世相似,印度洋东西侧SST梯度增加[105]。以上证据说明该时期湿润的气候条件可能受到海道关闭引起的沃克环流增强带来的影响。该时期海道关闭不仅维持了东南亚的湿润气候,还可能强化了全球降温背景下其他地区的干旱化事件,例如相邻的澳大利亚西北部、非洲等地区(图7)。此外,在晚中新世全球海平面显著下降(图7d)以及板块持续俯冲的背景下[32,34],东南亚岛屿和陆地的面积增加也可能对晚中新世湿润气候条件的维持做出了一定贡献[110–111]。综上所述,渐新世-中新世东南亚水文气候环境演变应该同时受到了区域构造运动和全球环境驱动因子的影响。
5. 结论与展望
(1) 孢粉记录、煤层沉积、浅海碳酸盐堆积和生物地理演化等多种证据表明,东南亚水文气候在渐新世与中新世之交发生了重大调整,从较为干旱状态转变为湿润条件,并且整个中新世都保持了较为湿润的气候条件。
(2)最近的模拟工作表明印尼海道关闭在东南亚干湿气候格局调整中可能扮演了重要角色。海道关闭可以通过影响次表层水体促使太平洋沃克环流增强,使东南亚由干转湿,并抵消了晚中新世全球降温的影响,进而维持湿润的气候格局。
(3)目前的工作仍然有限,海道关闭的时间和位置约束还不准确。模型中较为粗略的海陆格局和地形设置也在很大程度上影响了模拟输出结果。此外,目前印太暖池和东南亚水文气候演变历史还缺乏精细的重建工作,已有的数据揭示这二者发生转型的时间并不一致。未来还需要推进记录重建和数值模拟两方面的工作,以便更准确评估构造运动与气候演变的关系。
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图 1 印太暖池现代海流与本文研究所引用的站位资料
a:印太暖池分布范围及其表层洋流,其中橙色实线代表暖流,蓝色实线代表寒流,绿色实线代表季风淡水输入(修改自文献[37–38]);b:东南亚代表性钻井孢粉剖面(黑色圆点, 修改自文献[39–43])和主要含煤盆地分布(陆上褐色区域, 修改自文献[44–45]),华莱士生物区及其界限(白色虚线, 修改自文献[46])。
Figure 1. Modern oceanography in the Indo-Pacific Warm Pool and sites cited in this study
a: Distribution of the Indo-Pacific Warm Pool (black lines) and surface oceanic circulation in the region. Warm currents (orange lines), cold currents (blue lines), and monsoonal freshwater plumes (green lines) are presented (modified after references [37–38]); b: location of palynological profiles from representative wells (black dots, modified after references [39–43]) and main coal-bearing basins of Southeast Asia (brown patches on land, modified after references [44–45]), the Wallace’s zoogeographical region and its boundaries are denoted (white dashed lines, modified after reference [46]).
图 2 全球气候变化和印尼海道构造演变
a:全球深海底栖有孔虫δ18O[30];b:全球平均地表温度相对变化[31];c:全球海平面相对变化[32];d:大气CO2浓度[31];e:基于TEX86重建的全球SST,红色代表中低纬度地区,蓝色代表高纬地区[33];f:印尼海道构造演变重建(修改自文献[34,38])。
Figure 2. Global climate change and tectonic evolution of the Indonesian Seaway
a: Global deep-sea benthic foraminiferal δ18O[30]; b: variation in global mean surface temperature relative to the present estimated from benthic δ18O[31]; c: variation of sea level relative to the present estimated from benthic δ18O[32]; d: atmospheric CO2 estimate[31]; e: global SST estimated from TEX86, where red line represents middle-low latitude region and blue line represents high latitude region[33]; f: tectonic reconstruction of the Indonesian Seaway (modified after references [34,38]).
图 6 渐新世至中新世东南亚水文气候变化与浅海碳酸盐沉积发育情况对比
a:基于孢粉组合重建的水文气候变化,b:赤道不同碳酸盐生物相的相对面积,c:赤道不同碳酸盐生物相相对面积的相对比例,d:碳酸盐台地数量,e:碳酸盐浅滩和建隆数量,f:格架礁的发育情况(修改自文献[47,84])。
Figure 6. Comparison of hydrological changes in the Oligocene to Miocene with the development of shallow-marine carbonate formations in Southeast Asia
a: Reconstructed hydrological changes based on palynological assemblages, b: the relative areas of equatorial carbonate of different carbonate biofacies, c: the relative proportion of the relative areas of equatorial carbonate to different carbonate biofacies, d: numbers of carbonate platforms, e: numbers of carbonate shoals/buildups, f: development of framework reefs(modified after Refs. [47,84]).
图 7 渐新世至中新世东南亚与其他区域的水文气候对比
a:基于孢粉组合[84]和模拟[46]重建的东南亚水文气候变化;b:印尼海道演变历史[38,34];c:全球深海底栖有孔虫δ18O[30];d:全球平均地表温度相对变化[31]和海平面相对变化[32];e:大气CO2浓度[31];f:基于TEX86重建的全球SST,红色为中低纬地区,蓝色为高纬地区[33];g:巴基斯坦大型食草哺乳动物牙釉质δ18O和δ13C[48];h:亚洲内陆旱生植物占比,其中绿色散点为Ephedra + Nitraria + Chenopodiaceae + Artemisia之和[49,108],橙色柱状为干草甸和沙漠景观植被占比之和[50];i:非洲叶蜡烷烃δ13C[51–52];j:全球晚中新世C4植被扩张开始时间[51];k:澳大利亚西北部IODP U1464钾元素含量,大于0.3为湿润,小于0.2为干旱[53]。
Figure 7. Comparison of hydrological changes in the Oligocene to Miocene in Southeast Asia with other regions
a: Reconstructed hydrological changes based on the palynological assemblages[84]and model stimulation[46]; b: the evolution history of the Indonesian Seaway[38,34]; c: global deep-sea benthic foraminiferal δ18O[30]; d: global mean surface temperature[31] and sea level change[32] relative to the present estimated from benthic δ18O; e: atmospheric CO2 estimate[31]; f: global SST estimated from TEX86, where red line represents middle-low latitude region and blue line represents high latitude region[33]; g: δ18O and δ13C of tooth enamel bioapatite for large herbivorous mammal in Pakistan[48]; h: xerophytic vegetation changes in Central Asia, where green dots represent the sum of Ephedra + Nitraria + Chenopodiaceae + Artemisia[49,108], and orange bars represent the sum of steppe and desert vegetation[50]; i: compilation of δ13C for plant wax n-alkanes in Africa[51–52]; j: the onset of C4 vegetation expansion in the Late Miocene across the globe[51]; k: potassium content of the IODP U1464 located in the northwestern Australia, with wet and dry conditions being represented by >0.3 and <0.2, respectively[53].
表 1 主要引用资料基本信息
Table 1 Basic information of main references cited for this study
站位名称 地点 替代性指标 气候指示意义 覆盖时间 地层方法 参考文献 Well C 东南亚越南南部 孢粉组合相对丰度 植被类型干湿变化 24.2~28.2 Ma 生物地层 [42] Well E 东南亚越南南部 孢粉组合相对丰度 植被类型干湿变化 20.9~23.4 Ma 生物地层 [42] Belut-3 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 33.7~34.8 Ma 生物地层 [39] Kambing-1 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 27.5~29.4 Ma 生物地层 [39] Kambing-1? 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 23.8~27.5 Ma 生物地层 [39] Kadal-1 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 21.2~23.8 Ma 生物地层 [39] Bergading Deep-3 东南亚马来半岛东部 孢粉组合相对丰度 植被类型干湿变化 7.7~16.9 Ma 生物地层 [43] Dengkis-1 东南亚婆罗洲北部 孢粉组合相对丰度 植被类型干湿变化 6.6~14.2 Ma 生物地层 [43] Kuda Laut-1 东南亚婆罗洲北部 孢粉组合相对丰度 植被类型干湿变化 3.5~23.1 Ma 生物地层 [43] Bukoh-1 东南亚婆罗洲北部 孢粉组合相对丰度 植被类型干湿变化 20.2~28.2 Ma 生物地层 [43] Well A 东南亚婆罗洲东部 孢粉组合相对丰度 植被类型干湿变化 0~8.5 Ma 生物地层 [40] Well B 东南亚婆罗洲东部 孢粉组合相对丰度 植被类型干湿变化 0~9.4 Ma 生物地层 [40] Well X 东南亚爪哇海东部 孢粉组合相对丰度 植被类型干湿变化 23~30 Ma 生物地层 [41] Well Y 东南亚爪哇海东部 孢粉组合相对丰度 植被类型干湿变化 23~30 Ma 生物地层 [41] 东南亚 煤层分布 降雨强度干湿变化 新生代 [44–45] 东南亚 浅海碳酸盐生物相相对面积 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐生物相面积相对占比 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐台地数量 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐浅滩/建隆数量 降雨强度干湿变化 新生代 [47] 东南亚 浅海碳酸盐格架礁分布 降雨强度干湿变化 新生代 [47] 东南亚 降水量数值模拟 降雨强度干湿变化 0~30 Ma [46] 巴基斯坦 食草动物牙釉质碳同位素 植被类型干湿变化 0~33 Ma [48] 巴基斯坦 食草动物牙釉质氧同位素 降雨强度干湿变化 0~33 Ma [48] 中国西北 孢粉组合相对丰度 植被类型干湿变化 0~45 Ma [49] 中国西北 孢粉组合相对丰度 植被类型干湿变化 0~39 Ma [50] 非洲东西沿海 有机质叶蜡烷烃碳同位素 植被类型干湿变化 0~23 Ma [51] 非洲埃塞俄比亚 有机质叶蜡烷烃碳同位素 植被类型干湿变化 渐新世-早中新世 [52] IODP U1464 澳大利亚西北 沉积物钾元素含量 降雨强度干湿变化 5~15 Ma 生物地层 [53] 全球 底栖有孔虫氧同位素 全球冰量底层海水温度 新生代 [30] 全球 底栖有孔虫氧同位素 全球平均温度 新生代 [31] 全球 多指标(例如浮游植物长链不饱和酮类
碳同位素、硼同位素等)大气CO2浓度 新生代 [31] 全球 底栖有孔虫氧同位素 全球海平面 0~40 Ma [32] 全球 有机质TEX86 SST 新生代 [33] -
[1] De Deckker P. The Indo-Pacific Warm Pool: critical to world oceanography and world climate[J]. Geoscience Letters, 2016, 3(1):20. doi: 10.1186/s40562-016-0054-3
[2] Webster P J. The role of hydrological processes in ocean‐atmosphere interactions[J]. Reviews of Geophysics, 1994, 32(4):427-476. doi: 10.1029/94RG01873
[3] Oppo D W, Rosenthal Y. The great Indo-Pacific communicator[J]. Science, 2010, 328(5985):1492-1494. doi: 10.1126/science.1187273
[4] Rodgers K B, Latif M, Legutke S. Sensitivity of equatorial Pacific and Indian Ocean watermasses to the position of the Indonesian Throughflow[J]. Geophysical Research Letters, 2000, 27(18):2941-2944. doi: 10.1029/1999GL002372
[5] Cane M A, Molnar P. Closing of the Indonesian seaway as a precursor to east African aridification around 3–4 million years ago[J]. Nature, 2001, 411(6834):157-162. doi: 10.1038/35075500
[6] Karas C, Nürnberg D, Gupta A K, et al. Mid-Pliocene climate change amplified by a switch in Indonesian subsurface throughflow[J]. Nature Geoscience, 2009, 2(6):434-438. doi: 10.1038/ngeo520
[7] Karas C, Nürnberg D, Tiedemann R, et al. Pliocene climate change of the Southwest Pacific and the impact of ocean gateways[J]. Earth and Planetary Science Letters, 2011, 301(1-2):117-124. doi: 10.1016/j.jpgl.2010.10.028
[8] Bali H, Gupta A K, Mohan K, et al. Evolution of the oligotrophic West Pacific Warm Pool during the Pliocene‐Pleistocene boundary[J]. Paleoceanography and Paleoclimatology, 2020, 35(11):e2020PA003875. doi: 10.1029/2020PA003875
[9] Hall R. Cenozoic geological and plate tectonic evolution of SE Asia and the SW Pacific: computer-based reconstructions, model and animations[J]. Journal of Asian Earth Sciences, 2002, 20(4):353-431. doi: 10.1016/S1367-9120(01)00069-4
[10] Kuhnt W, Holbourn A, Hall R, et al. Neogene history of the Indonesian Throughflow[M]//Clift P, Kuhnt W, Wang P, et al. Continent-Ocean Interactions within East Asian Marginal Seas. Washington: American Geophysical Union, 2004: 299-320.
[11] Li Q Y, Jian Z M, Su X. Late Oligocene rapid transformations in the South China Sea[J]. Marine Micropaleontology, 2005, 54(1-2):5-25. doi: 10.1016/j.marmicro.2004.09.008
[12] Woodruff F, Savin S M. Miocene deepwater oceanography[J]. Paleoceanography, 1989, 4(1):87-140. doi: 10.1029/PA004i001p00087
[13] Von Der Heydt A S, Dijkstra H A. The impact of ocean gateways on ENSO variability in the Miocene[M]//Hall R, Cottam M A, Wilson M E J. The SE Asian Gateway: History and Tectonics of the Australia-Asia Collision. London: Geological Society of London, 2011: 305-318.
[14] Li Q Y, Li B H, Zhong G F, et al. Late Miocene development of the western Pacific warm pool: planktonic foraminifer and oxygen isotopic evidence[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2006, 237(2-4):465-482. doi: 10.1016/j.palaeo.2005.12.019
[15] Jian Z M, Yu Y Q, Li B H, et al. Phased evolution of the south–north hydrographic gradient in the South China Sea since the middle Miocene[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2006, 230(3-4):251-263. doi: 10.1016/j.palaeo.2005.07.018
[16] Nathan S A, Leckie R M. Early history of the Western Pacific Warm Pool during the middle to late Miocene (~13.2–5.8 Ma): role of sea-level change and implications for equatorial circulation[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2009, 274(3-4):140-159. doi: 10.1016/j.palaeo.2009.01.007
[17] Gourlan A T, Meynadier L, Allègre C J. Tectonically driven changes in the Indian Ocean circulation over the last 25 Ma: neodymium isotope evidence[J]. Earth and Planetary Science Letters, 2008, 267(1-2):353-364. doi: 10.1016/j.jpgl.2007.11.054
[18] Sosdian S M, Lear C H. Initiation of the western Pacific warm pool at the middle Miocene climate transition?[J]. Paleoceanography and Paleoclimatology, 2020, 35(12):e2020PA003920. doi: 10.1029/2020PA003920
[19] Sato K, Oda M, Chiyonobu S, et al. Establishment of the western Pacific warm pool during the Pliocene: evidence from planktic foraminifera, oxygen isotopes, and Mg/Ca ratios[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2008, 265(1-2):140-147. doi: 10.1016/j.palaeo.2008.05.003
[20] Auer G, De Vleeschouwer D, Smith R A, et al. Timing and pacing of Indonesian Throughflow restriction and its connection to Late Pliocene Climate Shifts[J]. Paleoceanography and Paleoclimatology, 2019, 34(4):635-657. doi: 10.1029/2018PA003512
[21] 周祖翼, 金性春, 王嘹亮, 等. 印尼海道的两度关闭与西太平洋暖池的形成和兴衰[J]. 海洋地质与第四纪地质, 2004, 24(1):7-14 ZHOU Zuyi, JIN Xinchun, WANG Liaoliang, et al. Two closures of the Indonesian seaway and its relationship to the formation and evolution of the western Pacific warm pool[J]. Marine Geology & Quaternary Geology, 2004, 24(1):7-14.]
[22] Bayon G, Patriat M, Godderis Y, et al. Accelerated mafic weathering in Southeast Asia linked to late Neogene cooling[J]. Science Advances, 2023, 9(13):eadf3141. doi: 10.1126/sciadv.adf3141
[23] Tan N, Li H, Zhang Z S, et al. Closure of tropical seaways favors the climate and vegetation in tropical Africa and South America approaching their present conditions[J]. Global and Planetary Change, 2024, 233:104351. doi: 10.1016/j.gloplacha.2023.104351
[24] Brierley C M, Fedorov A V. Relative importance of meridional and zonal sea surface temperature gradients for the onset of the ice ages and Pliocene-Pleistocene climate evolution[J]. Paleoceanography, 2010, 25(2):PA2214.
[25] Fedorov A V, Burls N J, Lawrence K T, et al. Tightly linked zonal and meridional sea surface temperature gradients over the past five million years[J]. Nature Geoscience, 2015, 8(12):975-980. doi: 10.1038/ngeo2577
[26] Su Q D, Nie J S, Meng Q Q, et al. Central Asian drying at 3.3 Ma linked to tropical forcing?[J]. Geophysical Research Letters, 2019, 46(17-18):10561-10567. doi: 10.1029/2019GL084648
[27] 谭宁, 张仲石, 郭正堂, 等. 上新世热带海道变化影响东亚气候的模拟研究[J]. 地学前缘, 2022, 29(5):310-321 TAN Ning, ZHANG Zhongshi, GUO Zhengtang, et al. Modeling study of the impact of tropical seaway changes on East Asian climate[J]. Earth Science Frontiers, 2022, 29(5):310-321.]
[28] Krebs U, Park W, Schneider B. Pliocene aridification of Australia caused by tectonically induced weakening of the Indonesian throughflow[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2011, 309(1-2):111-117. doi: 10.1016/j.palaeo.2011.06.002
[29] Christensen B A, Renema W, Henderiks J, et al. Indonesian throughflow drove Australian climate from humid Pliocene to arid Pleistocene[J]. Geophysical Research Letters, 2017, 44(13):6914-6925. doi: 10.1002/2017GL072977
[30] Westerhold T, Marwan N, Drury A J, et al. An astronomically dated record of Earth’s climate and its predictability over the last 66 million years[J]. Science, 2020, 369(6509):1383-1387. doi: 10.1126/science.aba6853
[31] The Cenozoic CO2 Proxy Integration Project (CenCO2PIP) Consortium. Toward a Cenozoic history of atmospheric CO2[J]. Science, 2023, 382(6675):eadi5177.
[32] Rohling E J, Yu J M, Heslop D, et al. Sea level and deep-sea temperature reconstructions suggest quasi-stable states and critical transitions over the past 40 million years[J]. Science Advances, 2021, 7(26):eabf5326. doi: 10.1126/sciadv.abf5326
[33] Auderset A, Moretti S, Taphorn B, et al. Enhanced ocean oxygenation during Cenozoic warm periods[J]. Nature, 2022, 609(7925):77-82. doi: 10.1038/s41586-022-05017-0
[34] Hall R. Late Jurassic–Cenozoic reconstructions of the Indonesian region and the Indian Ocean[J]. Tectonophysics, 2012, 570-571:1-41. doi: 10.1016/j.tecto.2012.04.021
[35] Gaina C, Müller D. Cenozoic tectonic and depth/age evolution of the Indonesian gateway and associated back-arc basins[J]. Earth-Science Reviews, 2007, 83(3-4):177-203. doi: 10.1016/j.earscirev.2007.04.004
[36] Hall R. Australia–SE Asia collision: plate tectonics and crustal flow[M]//Hall R, Cottam M A, Wilson M E J. The SE Asian Gateway: History and Tectonics of the Australia-Asia Collision. London: Geological Society of London, 2011: 75-109.
[37] Petrick B, Martínez-García A, Auer G, et al. Glacial Indonesian throughflow weakening across the mid-Pleistocene climatic transition[J]. Scientific Reports, 2019, 9(1):16995. doi: 10.1038/s41598-019-53382-0
[38] Gallagher S J, Auer G, Brierley C M, et al. Cenozoic history of the Indonesian gateway[J]. Annual Review of Earth and Planetary Sciences, 2024, 52:581-604. doi: 10.1146/annurev-earth-040722-111322
[39] Morley R J, Morley H P, Restrepo-Pace P. Unravelling the tectonically controlled stratigraphy of the West Natuna Basin by means of palaeo-derived Mid Tertiary climate changes[C]//Proceedings of the 29th Annual Convention Proceedings. AAPG, 2003: 1-24.
[40] Morley R J, Morley H P. Neogene climate history of the Makassar Straits, with emphasis on the Attaka Region, East Kalimantan, Indonesia[C]//Proceedings of the 34th Annual Convention Proceedings. AAPG, 2010.
[41] Lelono E B, Morley R J. Oligocene palynological zonation scheme from East Java Sea[J]. Scientific Contributions Oil & Gas, 2011, 34(2):95-104.
[42] Morley R J, Dung B V, Tung N T, et al. High-resolution Palaeogene sequence stratigraphic framework for the Cuu Long Basin, offshore Vietnam, driven by climate change and tectonics, established from sequence biostratigraphy[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2019, 530:113-135. doi: 10.1016/j.palaeo.2019.05.010
[43] Morley R J, Hasan S S, Morley H P, et al. Sequence biostratigraphic framework for the Oligocene to Pliocene of Malaysia: high-frequency depositional cycles driven by polar glaciation[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2021, 561:110058. doi: 10.1016/j.palaeo.2020.110058
[44] Davis R C, Noon S W, Harrington J. The petroleum potential of Tertiary coals from Western Indonesia: relationship to mire type and sequence stratigraphic setting[J]. International Journal of Coal Geology, 2007, 70(1-3):35-52. doi: 10.1016/j.coal.2006.02.008
[45] Friederich M C, Moore T A, Flores R M. A regional review and new insights into SE Asian Cenozoic coal-bearing sediments: why does Indonesia have such extensive coal deposits?[J]. International Journal of Coal Geology, 2016, 166:2-35. doi: 10.1016/j.coal.2016.06.013
[46] Skeels A, Boschman L M, Mcfadden I R, et al. Paleoenvironments shaped the exchange of terrestrial vertebrates across Wallace’s Line[J]. Science, 2023, 381(6653):86-92. doi: 10.1126/science.adf7122
[47] Wilson M E J. Global and regional influences on equatorial shallow-marine carbonates during the Cenozoic[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2008, 265(3-4):262-274. doi: 10.1016/j.palaeo.2008.05.012
[48] Martin C, Bentaleb I, Antoine P O. Pakistan mammal tooth stable isotopes show paleoclimatic and paleoenvironmental changes since the early Oligocene[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2011, 311(1-2):19-29. doi: 10.1016/j.palaeo.2011.07.010
[49] Barbolini N, Woutersen A, Dupont-Nivet G, et al. Cenozoic evolution of the steppe-desert biome in Central Asia[J]. Science Advances, 2020, 6(41):eabb8227. doi: 10.1126/sciadv.abb8227
[50] Jia Y X, Wu H B, Zhu S Y, et al. Cenozoic aridification in Northwest China evidenced by paleovegetation evolution[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2020, 557:109907. doi: 10.1016/j.palaeo.2020.109907
[51] Polissar P J, Rose C, Uno K T, et al. Synchronous rise of African C4 ecosystems 10 million years ago in the absence of aridification[J]. Nature Geoscience, 2019, 12(8):657-660. doi: 10.1038/s41561-019-0399-2
[52] Krawielicki J. Coupled climate, ecosystem and landscape development in the Afro-Mediterranean region since the Oligocene[D]. Doctor Dissertation of ETH Zürich, 2019.
[53] Groeneveld J, Henderiks J, Renema W, et al. Australian shelf sediments reveal shifts in Miocene Southern Hemisphere westerlies[J]. Science Advances, 2017, 3(5):e1602567. doi: 10.1126/sciadv.1602567
[54] Sia S G, Abdullah W H, Konjing Z, et al. The age, palaeoclimate, palaeovegetation, coal seam architecture/mire types, paleodepositional environments and thermal maturity of syn-collision paralic coal from Mukah, Sarawak, Malaysia[J]. Journal of Asian Earth Sciences, 2014, 81:1-19. doi: 10.1016/j.jseaes.2013.11.014
[55] Sia S G, Abdullah W H, Konjing Z, et al. Floristic and climatic changes at the Balingian Province of the Sarawak Basin, Malaysia, in response to Neogene global cooling, aridification and grassland expansion[J]. CATENA, 2019, 173:445-455. doi: 10.1016/j.catena.2018.10.044
[56] Konjing Z, Abd Rahman A H, Ismail M S, et al. Late Oligocene-Early Miocene palynological succession from marginal marine deposits, Nyalau Formation, Bintulu Sarawak: palynostratigraphy, paleovegetation and paleoclimate significance[J]. Bulletin of the Geological Society of Malaysia, 2022, 74:17-41. doi: 10.7186/bgsm74202202
[57] Lelono E B. Pollen records from the Oligocene of western Indonesia as the evidences of climate changes[J]. Scientific Contributions Oil & Gas, 2017, 40(3):107-115.
[58] Bao X J, Hu Y Y, Scotese C R, et al. Quantifying climate conditions for the formation of coals and evaporites[J]. National Science Review, 2023, 10(6):nwad051. doi: 10.1093/nsr/nwad051
[59] Flores R M. Origin of coal as gas source and reservoir rocks[M]//Flores R M. Coal and Coalbed Gas. Amsterdam: Elsevier, 2014: 97-165.
[60] 陈槐, 吴宁, 王艳芬, 等. 泥炭沼泽湿地研究的若干基本问题与研究简史[J]. 中国科学: 地球科学, 2021, 51(1):15-26 doi: 10.1360/SSTe-2020-0073 CHEN Huai, WU Ning, WANG Yanfen, et al. A historical overview about basic issues and studies of mires[J]. Scientia Sinica Terrae, 2021, 51(1):15-26.] doi: 10.1360/SSTe-2020-0073
[61] Stock A T, Littke R, Lücke A, et al. Miocene depositional environment and climate in western Europe: the lignite deposits of the Lower Rhine Basin, Germany[J]. International Journal of Coal Geology, 2016, 157:2-18. doi: 10.1016/j.coal.2015.06.009
[62] Guo Q L, Littke R, Zieger L. Petrographical and geochemical characterization of sub-bituminous coals from mines in the Cesar-Ranchería Basin, Colombia[J]. International Journal of Coal Geology, 2018, 191:66-79. doi: 10.1016/j.coal.2018.03.008
[63] Petersen H I, Andersen C, Anh P H, et al. Petroleum potential of Oligocene lacustrine mudstones and coals at Dong Ho, Vietnam: an outcrop analogue to terrestrial source rocks in the greater Song Hong Basin[J]. Journal of Asian Earth Sciences, 2001, 19(1-2):135-154. doi: 10.1016/S1367-9120(00)00022-5
[64] Susilawati R. Minerals and inorganic matter in coals of the Bukit Asam Coalfield, South Sumatra Basin, Indonesia[D]. Master Dissertation of University of New South Wales, 2004.
[65] Petersen H I, Foopatthanakamol A, Ratanasthien B. Petroleum potential, thermal maturity and the oil window of oil shales and coals in Cenozoic rift basins, central and Northern Thailand[J]. Journal of Petroleum Geology, 2006, 29(4):337-360. doi: 10.1111/j.1747-5457.2006.00337.x
[66] Petersen H I, Lindström S, Nytoft H P, et al. Composition, peat-forming vegetation and kerogen paraffinicity of Cenozoic coals: relationship to variations in the petroleum generation potential (Hydrogen Index)[J]. International Journal of Coal Geology, 2009, 78(2):119-134. doi: 10.1016/j.coal.2008.11.003
[67] Singh P K, Singh M P, Singh A K, et al. Petrographic characteristics of coal from the Lati Formation, Tarakan basin, East Kalimantan, Indonesia[J]. International Journal of Coal Geology, 2010, 81(2):109-116. doi: 10.1016/j.coal.2009.11.006
[68] Widodo S, Oschmann W, Bechtel A, et al. Distribution of sulfur and pyrite in coal seams from Kutai Basin (East Kalimantan, Indonesia): implications for paleoenvironmental conditions[J]. International Journal of Coal Geology, 2010, 81(3):151-162. doi: 10.1016/j.coal.2009.12.003
[69] Alias F L, Abdullah W H, Hakimi M H, et al. Organic geochemical characteristics and depositional environment of the Tertiary Tanjong Formation coals in the Pinangah area, onshore Sabah, Malaysia[J]. International Journal of Coal Geology, 2012, 104:9-21. doi: 10.1016/j.coal.2012.09.005
[70] Hakimi M H, Abdullah W H, Alias F L, et al. Organic petrographic characteristics of Tertiary (Oligocene–Miocene) coals from eastern Malaysia: rank and evidence for petroleum generation[J]. International Journal of Coal Geology, 2013, 120:71-81. doi: 10.1016/j.coal.2013.10.003
[71] Chaiseanwang P, Chenrai P. Organic geochemical characteristics of Mae Teep coal deposits, Thailand[J]. ScienceAsia, 2020, 46S:102-109. doi: 10.2306/scienceasia1513-1874.2020.S015
[72] Sattraburut T, Ratanasthien B, Thasod Y. Palaeovegetation and palaeoclimate of tertiary sediments from Hongsa Coalfield, Xayabouly province, Lao PDR: implication from palynofloras[J]. Songklanakarin Journal of Science and Technology, 2021, 43(3):648-659.
[73] Fikri H N, Sachsenhofer R F, Bechtel A, et al. Organic geochemistry and petrography in Miocene coals in the Barito Basin (Tutupan Mine, Indonesia): evidence for astronomic forcing in kerapah type peats[J]. International Journal of Coal Geology, 2022, 256:103997. doi: 10.1016/j.coal.2022.103997
[74] Petersen H I, Fyhn M B W, Nytoft H P, et al. Miocene coals in the Hanoi Trough, onshore northern Vietnam: depositional environment, vegetation, maturity, and source rock quality[J]. International Journal of Coal Geology, 2022, 253:103953. doi: 10.1016/j.coal.2022.103953
[75] Patria A A, Suhendra R, Anggara F, et al. Association and textural-compositional evolution of pyrite-organic matter in coals of the Tarakan, Barito, and Pasir Basins, Kalimantan, Indonesia[J]. International Journal of Coal Geology, 2024, 282:104442. doi: 10.1016/j.coal.2023.104442
[76] Dai S F, Bechtel A, Eble C F, et al. Recognition of peat depositional environments in coal: a review[J]. International Journal of Coal Geology, 2020, 219:103383. doi: 10.1016/j.coal.2019.103383
[77] Sun Y Z. Review and update on the applications of inertinite macerals in coal geology, paleoclimatology, and paleoecology[J]. Palaeoworld, 2024.
[78] Ruan Y, Mohtadi M, Dupont L M, et al. Interaction of fire, vegetation, and climate in tropical ecosystems: a multiproxy study over the past 22, 000 years[J]. Global Biogeochemical Cycles, 2020, 34(11):e2020GB006677. doi: 10.1029/2020GB006677
[79] Sen S, Naskar S, Das S. Discussion on the concepts in paleoenvironmental reconstruction from coal macerals and petrographic indices[J]. Marine and Petroleum Geology, 2016, 73:371-391. doi: 10.1016/j.marpetgeo.2016.03.015
[80] Wilson M E J, Vecsei A. The apparent paradox of abundant foramol facies in low latitudes: their environmental significance and effect on platform development[J]. Earth-Science Reviews, 2005, 69(1-2):133-168. doi: 10.1016/j.earscirev.2004.08.003
[81] Wilson M E J. Equatorial carbonates: an earth systems approach[J]. Sedimentology, 2012, 59(1):1-31. doi: 10.1111/j.1365-3091.2011.01293.x
[82] Reijmer J J G. Marine carbonate factories: review and update[J]. Sedimentology, 2021, 68(5):1729-1796. doi: 10.1111/sed.12878
[83] Wilson M E J. Cenozoic carbonates in Southeast Asia: implications for equatorial carbonate development[J]. Sedimentary Geology, 2002, 147(3-4):295-428. doi: 10.1016/S0037-0738(01)00228-7
[84] Morley R J. Assembly and division of the South and South-East Asian flora in relation to tectonics and climate change[J]. Journal of Tropical Ecology, 2018, 34(4):209-234. doi: 10.1017/S0266467418000202
[85] Morley R J. Cretaceous and Tertiary climate change and the past distribution of megathermal rainforests[M]//Bush M, Flenley J, Gosling W. Tropical Rainforest Responses to Climatic Change. Berlin: Springer, 2011: 1-34.
[86] Bansal M, Morley R J, Nagaraju S, et al. Southeast Asian Dipterocarp origin and diversification driven by Africa-India floristic interchange[J]. Science, 2022, 375(6579):455-460. doi: 10.1126/science.abk2177
[87] Klaus S, Morley R J, Plath M, et al. Biotic interchange between the Indian subcontinent and mainland Asia through time[J]. Nature Communications, 2016, 7(1):12132. doi: 10.1038/ncomms12132
[88] Joyce E M, Thiele K R, Slik J W F, et al. Plants will cross the lines: climate and available land mass are the major determinants of phytogeographical patterns in the Sunda–Sahul Convergence Zone[J]. Biological Journal of the Linnean Society, 2021, 132(2):374-387. doi: 10.1093/biolinnean/blaa194
[89] Li J T, Li Y, Klaus S, et al. Diversification of rhacophorid frogs provides evidence for accelerated faunal exchange between India and Eurasia during the Oligocene[J]. Proceedings of the National Academy of Sciences of the United States of America, 2013, 110(9):3441-3446.
[90] Chen J M, Prendini E, Wu Y H, et al. An integrative phylogenomic approach illuminates the evolutionary history of Old World tree frogs (Anura: Rhacophoridae)[J]. Molecular Phylogenetics and Evolution, 2020, 145:106724. doi: 10.1016/j.ympev.2019.106724
[91] Yuan L M, Deng X L, Jiang D C, et al. Geographical range evolution of the genus Polypedates (Anura: Rhacophoridae) from the Oligocene to present[J]. Zoological Research, 2021, 42(1):116-123. doi: 10.24272/j.issn.2095-8137.2020.246
[92] Zhu R X, Wang H J, Wang H J, et al. Multi-spherical interactions and mechanisms of hydrocarbon enrichment in the Southeast Asian archipelagic tectonic system[J]. Science China Earth Sciences, 2024, 67(2):566-583. doi: 10.1007/s11430-023-1254-4
[93] Sun X J, Wang P X. How old is the Asian monsoon system? palaeobotanical records from China[J]. Palaeogeography, Palaeoclimatology, Palaeoecology, 2005, 222(3-4):181-222. doi: 10.1016/j.palaeo.2005.03.005
[94] Guo Z T, Sun B, Zhang Z S, et al. A major reorganization of Asian climate by the early Miocene[J]. Climate of the Past, 2008, 4(3):153-174. doi: 10.5194/cp-4-153-2008
[95] Rea D K, Leinen M, Janecek T R. Geologic approach to the long-term history of atmospheric circulation[J]. Science, 1985, 227(4688):721-725. doi: 10.1126/science.227.4688.721
[96] Rea D K, Snoeckx H, Joseph L H. Late Cenozoic Eolian deposition in the North Pacific: Asian drying, Tibetan uplift, and cooling of the northern hemisphere[J]. Paleoceanography, 1998, 13(3):215-224. doi: 10.1029/98PA00123
[97] Dupont L M, Rommerskirchen F, Mollenhauer G, et al. Miocene to Pliocene changes in South African hydrology and vegetation in relation to the expansion of C4 plants[J]. Earth and Planetary Science Letters, 2013, 375:408-417. doi: 10.1016/j.jpgl.2013.06.005
[98] Polissar P J, Uno K T, Phelps S R, et al. Hydrologic changes drove the Late Miocene expansion of C4 grasslands on the Northern Indian subcontinent[J]. Paleoceanography and Paleoclimatology, 2021, 36(4):e2020PA004108. doi: 10.1029/2020PA004108
[99] Shen X Y, Wan S M, Colin C, et al. Increased seasonality and aridity drove the C4 plant expansion in Central Asia since the Miocene–Pliocene boundary[J]. Earth and Planetary Science Letters, 2018, 502:74-83. doi: 10.1016/j.jpgl.2018.08.056
[100] Li M J, Wan S M, Colin C, et al. Expansion of C4 plants in South China and evolution of East Asian monsoon since 35 Ma: black carbon records in the northern South China Sea[J]. Global and Planetary Change, 2023, 223:104079. doi: 10.1016/j.gloplacha.2023.104079
[101] Carrapa B, Clementz M, Feng R. Ecological and hydroclimate responses to strengthening of the Hadley circulation in South America during the Late Miocene cooling[J]. Proceedings of the National Academy of Sciences of the United States of America, 2019, 116(20):9747-9752.
[102] Latorre C, Quade J, McIntosh W C. The expansion of C4 grasses and global change in the late Miocene: stable isotope evidence from the Americas[J]. Earth and Planetary Science Letters, 1997, 146(1-2):83-96. doi: 10.1016/S0012-821X(96)00231-2
[103] Hyland E G, Sheldon N D, Smith S Y, et al. Late Miocene rise and fall of C4 grasses in the western United States linked to aridification and uplift[J]. GSA Bulletin, 2019, 131(1-2):224-234. doi: 10.1130/B32009.1
[104] Andrae J W, McInerney F A, Polissar P J, et al. Initial expansion of C4 vegetation in Australia during the Late Pliocene[J]. Geophysical Research Letters, 2018, 45(10):4831-4840. doi: 10.1029/2018GL077833
[105] Zhang R, Liu Z H, Jiang D B, et al. Cenozoic Indo-Pacific warm pool controlled by both atmospheric CO2 and paleogeography[J]. Science Bulletin, 2024, 69(9):1323-1331. doi: 10.1016/j.scib.2024.02.028
[106] Scotese C R. Atlas of Earth History: Volume 1, Paleogeography[M]. Arlington: PALEOMAP Project, 2001.
[107] Yan Q, Korty R, Zhang Z S, et al. Large shift of the Pacific Walker Circulation across the Cenozoic[J]. National Science Review, 2021, 8(5):nwaa101. doi: 10.1093/nsr/nwaa101
[108] Wu F, Fang X M, Yang Y B, et al. Reorganization of Asian climate in relation to Tibetan Plateau uplift[J]. Nature Reviews Earth & Environment, 2022, 3(10):684-700.
[109] Matsui H, Nishi H, Kuroyanagi A, et al. Vertical thermal gradient history in the eastern equatorial Pacific during the early to middle Miocene: implications for the equatorial thermocline development[J]. Paleoceanography, 2017, 32(7):729-743. doi: 10.1002/2016PA003058
[110] Dayem K E, Noone D C, Molnar P. Tropical western Pacific warm pool and maritime continent precipitation rates and their contrasting relationships with the Walker Circulation[J]. Journal of Geophysical Research: Atmospheres, 2007, 112(D6):D06101.
[111] Molnar P, Cronin T W. Growth of the Maritime Continent and its possible contribution to recurring Ice Ages[J]. Paleoceanography, 2015, 30(3):196-225. doi: 10.1002/2014PA002752