江南造山带西段北侧边界厘定:来自黔东大坪捕虏花岗岩的证据

Demarcation of North Boundary for Western Jiangnan Orogen: Evidence from Granitic Xenolith in Daping Area, East Guizhou

摘要

幔源岩浆可作为获取地球深部物质信息的“岩石探针”.对黔东大坪晚奥陶世（449 Ma）钾镁煌斑岩中首次发现的花岗岩捕虏体展开了锆石U-Pb年代学、Lu-Hf同位素以及锆石微量元素分析.结果表明，捕虏花岗岩锆石U-Pb谐和年龄为（833±2.6） Ma（MSWD=1.3， n=26），其锆石ε_Hf（t）值和亏损地幔模式年龄（T_DM）分别为-11.4~-2.30和2 457~1 893 Ma，锆石REE、U、Th、Pb、Nb、Hf等微量元素表明该捕虏体为造山作用相关S型花岗岩.花岗岩捕虏体与梵净山地区出露的新元古代花岗岩在结晶年龄、锆石Hf同位素组成上一致，暗示二者可能在深部共同构成一巨大花岗岩基，该花岗岩基为厘定江南造山带西段北侧边界提供了重要的物质证据.提出江南造山带西段北侧与扬子地块的边界应以张家界‒贵阳断裂为界.

Abstract

Mantle-derived magmas can serve as a “lithoprobe” for acquiring information about deep Earth materials. In this study, a newly discovered granite xenolith from the Late Ordovician (449 Ma) lamproite at Daping, East Guizhou, was investigated. Zircon U-Pb geochronology, Lu-Hf isotope, and trace element analyses were conducted on the granite xenolith. The results indicate that the concordant zircon U-Pb age of the granite xenolith is (833±2.6) Ma (MSWD=1.3, n=26). The ε_Hf(t) values range from -11.4 to -2.30, and the depleted mantle model ages (T_DM) vary from 2 457 to 1 893 Ma. The trace element compositions of zircon, including REEs, U, Th, Pb, Nb, and Hf, suggest that the granite xenolith is an S-type granite related to orogenic processes. The similarity in crystallization ages and Hf isotopic compositions between this granite xenolith and the Neoproterozoic granites exposed in the Fanjingshan region implies that they may collectively form a large granitic batholith at depth. This batholith provides crucial evidence for delineating the northern boundary of the western segment of the Jiangnan Orogen. It is proposed that the boundary between the northern side of the western Jiangnan Orogen and the Yangtze Block should be defined by the Zhangjiajie-Guiyang fault.

Graphical abstract

关键词

扬子地块 / 新元古代 / 岩石探针 / 捕虏体 / 造山带边界 / 花岗岩基 / 同位素 / 地质年代学.

Key words

Yangtze Block / Neoproterozoic / lithoprobe / xenolith / boundary of orogen belt / granitic batholith / isotopes / geochronology

0 引言

造山带为地壳的生长和改造提供了重要记录（Windley， 1992； Collins， 2002； Cawood et al.， 2009，2016）.花岗岩是大陆地壳形成、生长和再造的重要信使（徐夕生等， 2020），是大陆地壳形成演化过程的直接产物（翟明国等， 2016）.造山带内往往发育大量与造山作用相关的大花岗岩基，而花岗岩基通常拥有非均质的源区以及复杂的岩浆过程（Li et al.， 2024），通过对这些花岗岩研究可确定造山带类型，且对确定造山带地体边界以及理解造山带的形成演化过程至关重要（Cawood et al.， 2013； 王孝磊等， 2017； Yao et al.， 2019； Zhu et al.， 2023a）.然而，一些古老造山带由于后期稳定沉积地层覆盖和构造运动的强烈改造，使其重要信息被覆盖或改造，给造山带研究带来了困难（Shu et al.， 2021）.幸运的是造山带内出露的幔源岩浆可为学者研究造山带研究提供一个新的视角.钾镁煌斑岩作为来源深部地幔（Boyd and Gurney， 1986），成分特殊的幔源岩浆（Pearson et al.， 2003； Pearson and Wittig， 2014），其所携带的深部捕虏信息可以作为研究地球深部的天然“岩石探针”.例如Gardiner et al.（2020）对晚期金伯利岩中捕虏锆石研究揭示了北大西洋克拉通在太古代时期的形成演化以及Downes et al.（2007）对金伯利克拉通中部820 Ma金伯利岩所捕虏1 851 Ma花岗岩研究证实在古元古代地幔和地壳受到热事件和岩浆事件影响.此外岩石探针和同位素填图也可作为示踪造山带深部物源（物质）特征的重要手段，以揭示造山带不同地块在深部的物质组成和结构（侯增谦和王涛， 2018； 王涛等， 2022）.

江南造山带是位于扬子地块与华夏地块之间、形成于1 000~820 Ma的新元古代造山带（Zhao， 2015； 王孝磊等， 2017； Yao et al.， 2019； Shu et al.， 2021； Wang et al.， 2022），其成因模式目前仍然存有较大争议.Zhao（2015）认为江南造山带为华夏与扬子地块之间大洋双俯冲而形成陆‒陆软碰撞造山带，而后有诸多学者提出江南造山带为新元古代华夏对扬子地块俯冲‒增生所形成的增生造山带（王孝磊等， 2017； Yao et al.， 2019； Shu et al.， 2021）.近来又有学者认为江南造山带的形成过程可能更为复杂，提出为大怀玉地体与扬子地块俯冲碰撞而形成江南造山带，扬子与华夏地块的最终拼合可能在早古生代晚期（Wang et al.， 2022）.Yao et al. （2019）根据江南造山带空间分布将其划分为东段（怀玉或双溪坞地体）、中段（九岭地体）和西段三个部分.江南造山带东段和中段由于蛇绿岩套以及弧火山岩的出露（王孝磊等， 2017； Shu et al.， 2021），已有丰富的地球化学（Yao et al.， 2016； Yu et al.， 2016）和地球物理（He et al.， 2013； Dong et al.， 2015； Guo and Gao， 2018； 陈昌昕等， 2022； 严加永等， 2022； Han et al.， 2023）研究，基本确定了江南造山带东段和中段与扬子地块和华夏地块的界线，其北部界线为九江‒石台‒大庸（张家界）断裂，其南部界线为绍兴‒江山‒萍乡断裂.遗憾的是，造山带西段同期岩石多被<820 Ma的沉积地层所覆盖，导致其西段边界模糊不清，仅有少量地球物理研究对江南造山带西段北侧界线进行大致厘定（Dong et al.， 2015； Guo and Gao， 2018； 严加永等， 2022； 李雪垒等， 2023）.从Dong et al.（2015）对江南造山带中段地震解译可以看出在大庸断裂以南基底地层广泛发育大型褶皱冲断带，表明大庸断裂为江南造山带中段与扬子地块的边界，但继续往西的延伸则争议较大.有学者认为界线应过吉首‒铜仁‒凯里一线（严加永等， 2022），也有学者认为界线应为大庸‒铜仁‒河池‒百色一线（Guo and Gao， 2018； 李雪垒等， 2023），但二者在区域上应该为河池‒张家界一线.戴传固等（2013）对贵州大型隐伏断裂研究认为木黄‒贵阳‒普安断裂为江南造山带西侧北缘边界.以上对于江南造山带西段边界认识存在较大争议，因此准确厘定江南造山带西侧北缘边界需要更为直接的物质信息佐证，而这些信息目前大部分深埋地下.

庆幸的是，江南造山带西段存在早古生代金伯利岩‒钾镁煌斑岩等岩浆活动（Zhang et al.， 2023a），有望将深埋地下的江南造山带物质信息带至地表.据报道，黔东施秉大坪地区的钾镁煌斑岩中就包含有大量捕虏体（杨光忠等， 2019）.王亮等（2020）通过地球物理研究提出在大坪地区深部可能存在隐伏花岗岩.因此，大坪的钾镁煌斑岩存在捕虏江南造山带古老物质的可能性.鉴于此，笔者团队对在大坪岩体以往所打取岩心进行了细致观察，首次发现该岩体存在花岗岩捕虏体.本文进而以大坪钾镁煌斑岩中所捕虏的花岗岩为研究对象，对花岗岩捕虏体开展了锆石LA⁃ICP⁃MS U⁃Pb同位素年代学研究、锆石微量元素以及锆石LA⁃MC⁃ICP⁃MS Lu⁃Hf同位素分析与研究，为限定江南造山带西段北侧边界研究提供深部物质地球化学证据支撑.

1 地质背景

1.1　华南地质演化

华南板块主要由三大地质块体组成，分别为北西侧的扬子地块和南东侧华夏地块以及二者之间的江南造山带（Zhao and Cawood， 2012； Zheng et al.， 2013； Shu et al.， 2021）（图1）.其中，扬子地块是具有3.5 Ga陆核的古老克拉通（Zhang et al.， 2006； Zhao et al.， 2021）和一个广泛未暴露的太古宙基底，并经历多期地壳生长（Zheng et al.， 2006； Xiang et al.， 2018，2022）.华夏地块则主要由古元古代基底岩石（Xu et al.， 2007； Yu et al.， 2009）以及新元古代到早古生代沉积地层所构成（Zhao and Cawood， 2012； Shu et al.， 2021）.介于二者之间的江南造山带为1 000~820 Ma的新元古代造山带（Zhao， 2015； 王孝磊等， 2017； Yao et al.， 2019； Shu et al.， 2021； Wang et al.， 2022），并在造山作用结束后形成810~760 Ma的南华裂谷盆地和双峰式岩浆作用.在760~460 Ma之间，华南地块处于稳定的滨浅海‒斜坡沉积环境（Shu et al.， 2021）.随后，华南在原扬子板块区域（江南造山带西段）在早古生代发育了488 Ma和450 Ma两期金伯利岩‒钾镁煌斑岩岩浆活动，有超730多条大小岩脉在2 600 km²范围集中出露（Zhang et al.， 2023a），一些岩体还有大量捕虏体（杨光忠等， 2019）.

1.2　江南造山带新元古代花岗岩

江南造山带中广泛存在新元古代时期花岗岩岩浆作用，自东往西在江南造山带东段主要分布有休宁、歙县、许村岩体，中段主要以九岭花岗岩基、蓝田花岗岩、张邦源岩体等为代表，而西段则有梵净山花岗岩、桂北三防岩体、元宝山岩体等出露.江南造山带内最古老花岗岩为赣西北蛇绿岩带中的斜长花岗岩包体，其锆石的加权平均²⁰⁶Pb/²³⁸U结晶年龄为（970±21） Ma，被认为是由扬子地块南缘新元古代（970 Ma）岛弧下地幔楔形成的玄武岩质母岩浆演化而来（Gao et al.， 2009）.造山带中出露花岗岩主要以S型花岗岩为主（李献华等， 2001），但也有少量的I型（马铁球等， 2009； Liu and Zhao， 2018）、A型花岗岩（Huang et al.， 2018）报道.S型花岗岩主要分布于江南造山带中段的九岭岩体以及江南造山带西段的梵净山花岗岩、桂北三防岩体、元宝山岩体等，岩体侵位年龄主要集中在860~ 825 Ma（王孝磊等， 2017； 黄思访， 2021）.I型主要为形成于（821±6） Ma和（816±5） Ma的湖南张邦源岩体和蓝田岩体（马铁球等， 2009； Liu and Zhao， 2018）.

1.3　研究区地质概况

黔东大坪钾镁煌斑岩位于江南造山带西段梵净山南西约100 km（图1）.研究区出露地层从老到新主要有寒武系第二统耙榔组、清虚洞组一段和二段，寒武系第三统高台组、石冷水组.在大坪地区发现2条钾镁煌斑岩脉（图2b），均产出于寒武系第二统耙榔组粉砂质泥（页）岩.靠近江凯村的岩脉厚0.5 m，呈东西向横穿河道近直立产出，出露长度大于200 m.岩石主要由金云母、蛇纹石、绿泥石、白云石、石英、含钛矿物、褐铁矿等组成，蚀变强烈，以富含鲕粒胶黄铁矿为典型特征（杨光忠等， 2019）.靠近大坪的岩体地表出露宽度大于10 m，长度大于15 m，山体斜坡下部200 m范围内普见其转块，体积为0.2~0.5 m³，最大约5 m³，钾镁煌斑岩含大量围岩角砾捕虏体（25%~30%）（杨光忠等， 2019）（图2a~2c），该岩体侵位于晚奥陶世，岩浆磷灰石U⁃Pb年龄为（449±12） Ma （MSWD=2.3， n=38）（Zhang et al.， 2023a）.

1.4　样品采集

本研究对大坪钾镁煌斑岩岩体（GPS坐标： 108°12′46″ E， 27°5′43″ N）开展了详细的野外地质调查，由于岩体抗风化能力较弱，长期暴露地表处已风化为土黄色、黄褐色等土状，新鲜面则为灰色致密块状（图2a~2d）.岩体中可见大量捕虏体（图2a~2c）.前期工作有地勘队伍对该岩体进行了钻探工作，本次工作对钻探工作获取的新鲜岩心进行了详细观察，发现了与岩体存在显著差异的花岗岩捕虏体.花岗岩呈浅灰色，中细粒半自形‒他形、二长结构，块状构造（图2e~2f），主要矿物成分由斜长石（约35%）、碱性长石（约34%）、石英（约25%）组成，次要矿物由黑云母（约5%）以及其他矿物（约1%）组成.斜长石呈半自形‒他形板状结构，聚片双晶发育，粒径约1.5~3 mm.碱性长石呈现出他形结构，表面模糊，具有较强的高岭石化，粒径约0.5~ 1.5 mm.石英多呈半自形‒他形粒状结构，部分石英可见蠕英结构，粒径约0.5~3 mm.黑云母多呈片状，蚀变强烈.捕虏花岗岩最终定名为二长花岗岩.

2 分析方法

2.1　捕虏花岗岩锆石U⁃Pb年代学和微量元素分析

将捕虏花岗岩样品（20DP）去除明显风化部分后破碎至80目，利用磁性和重液分选方法对其中矿物进行分离，然后通过手工挑选锆石颗粒，并安置于环氧树脂盘上.利用光学方法（反射和透射光显微镜）、扫描电镜和阴极发光显微镜检查锆石颗粒，选择合适部位进行测试.锆石U⁃Pb同位素定年在北京燕都中实测试技术有限公司使用LA⁃ICP⁃MS完成.锆石剥蚀直径为32 μm，频率为8 Hz.激光剥蚀过程中采用氦气作载气、氩气为补偿气以调节灵敏度，二者在进入ICP之前通过一个Y型接头混合.每个时间分辨分析数据包括30 s的空白信号和50 s的样品信号.U⁃Pb 同位素定年中采用锆石标准Plesovice（Sláma et al.， 2008）作外标进行同位素分馏校正，并对91500（Wiedenbeck et al.， 1995）、Tanz（Hu et al.， 2021）、ZS（Ling et al.， 2022）分析一次作为监控.锆石微量元素含量利用SRM610作为多外标、Si作内标的方法进行定量计算（Liu et al.， 2010）.对分析数据的离线处理采用实验室自主开发的ZSkits软件完成.锆石样品的U⁃Pb年龄谐和图绘制和年龄权重平均计算均采用IsoplotR完成（Vermeesch， 2018）.

2.2　捕虏花岗岩锆石Lu⁃Hf同位素分析方法

锆石原位Lu⁃Hf同位素测试在中国科学院地球化学研究所使用美国热电公司（Thermo Fisher Scientific）Nepture⁃plus MC⁃ICP⁃MS与NWR193激光系统完成.测试步骤与校准方法参照Wu et al. （2006a）.锆石剥蚀使用频率为8 Hz，能量为16 J/cm²的激光剥蚀31 s，剥蚀出直径30 μm的剥蚀坑.测试时，由于锆石中的¹⁷⁶Lu/¹⁷⁷Hf比值极低（一般小于0.002），¹⁷⁶Lu对¹⁷⁶Hf的同位素干扰可以忽略不计.每个测试点的¹⁷²Yb/¹⁷³Yb平均值用于计算Yb的分馏系数，然后再扣除¹⁷⁶Yb对¹⁷⁶Hf的同质异位素干扰.¹⁷²Yb/¹⁷³Yb的同位素比值为1.352 74（Chu et al.， 2002）.样品同位素数据采用ICP⁃MS DataCal软件处理（Liu et al.， 2010）.

3 结果

3.1　捕虏花岗岩锆石U⁃Pb年龄

捕虏花岗岩U⁃Pb同位素比值及年龄见附表1.锆石颗粒大多呈长柱状，极少部分为短柱状或次圆状.锆石振荡环带清晰、整体棱角分明，自形程度较好，粒径介于60~120 μm（图3）.所测40颗锆石Th/U介于0.11~0.52，表明皆为岩浆锆石（Rubatto， 2002），这些锆石测点的²⁰⁷Pb/²⁰⁶Pb年龄范围在1 254~823 Ma.部分锆石样品存在一定的Pb丢失（20DP⁃08、20DP⁃09、20DP⁃16、20DP⁃21）（图4a），因此本文采用²⁰⁷Pb/²⁰⁶Pb年龄，这是由于样品在发生Pb丢失过程中，²⁰⁶Pb和²⁰⁷Pb将以同样的比例丢失，样品剩余的²⁰⁷Pb/²⁰⁶Pb和丢失前的比值一致，从而较大限度地消除Pb丢失对年龄测定的影响（张宏飞和高山， 2012）.其中26个测点（图4红色部分）²⁰⁷Pb/²⁰⁶Pb年龄为855~823 Ma，加权平均年龄为（836±6.6） Ma（MSWD=0.23， n=26）（图4b），与其U⁃Pb谐和年龄（833±2.6） Ma（MSWD=1.3， n=26）一致.

3.2　捕虏花岗岩锆石微量元素特征

捕虏花岗岩锆石微量元素含量分析结果见附表2.40颗捕虏花岗岩锆石∑REE变化范围为426×10^-6~3 065×10^-6（平均1 550×10^-6），轻稀土（LREE）为2×10^-6~384×10^-6（平均18×10^-6），重稀土（HREE）423×10^-6~3 056×10^-6（平均1 533×10^-6），LREE/HREE为0.002 2~0.187 7（平均0.010 6），具有LREE亏损和HREE明显富集的左倾配分模式.测点Ce/Ce*为0.95~20.83（平均5.18），Eu/Eu*为0.01~0.13（平均0.028），表现为显著的正Ce异常和负Eu异常.在40颗锆石测点中有3颗锆石（20DP⁃16、20DP⁃23、20DP⁃29）存在一定的LREE含量较高（图5a），但是这3颗锆石Th/U为0.25~0.41，表明为岩浆结晶锆石，可能为分析测试中锆石存在微小包裹体（苗壮等， 2020）.

3.3　捕虏花岗岩锆石Lu⁃Hf同位素

对捕虏花岗岩计入年龄统计的26颗锆石进行了Lu⁃Hf同位素分析，分析结果见附表3.所有锆石测点¹⁷⁶Yb/¹⁷⁷Hf介于0.008 487~0.065 870（平均0.035 437），¹⁷⁶Lu/¹⁷⁷Hf介于0.000 324~0.002 527（平均0.001 365），绝大多数锆石¹⁷⁶Lu/¹⁷⁷Hf比值远小于0.002（仅20DP⁃11、20DP⁃39分别为0.002 251、0.002 527），表明这些锆石在形成以后，不具有明显的放射成因Hf积累，因此所测¹⁷⁶Hf/¹⁷⁷Hf比值可以代表其形成锆石时体系的Hf同位素组成（吴福元等， 2007）.另外， 所有测试点的f_Lu/Hf值为-0.99到-0.92（平均-0.96），所有f_Lu/Hf值小于铁镁质地壳-0.34（Amelin et al.， 2000）和硅铝质地壳-0.72（Vervoort et al.， 1996），因此二阶段模式年龄更能真实反映源区物质在地壳的存留年龄.通过对所测结果计算的ε_Hf（t）值和亏损地幔模式年龄（T_DM）范围分别为-11.4~-2.3和2 457~1 893 Ma.

4 讨论

4.1　江南造山带西段约830 Ma隐伏S型花岗岩基

4.1.1　大坪捕虏花岗岩为S型花岗岩

由于所研究样品为钻孔所揭露，样品在制作薄片和分选单矿物后，不够测试全岩主微量含量，因而通过进一步分析捕虏体花岗岩的锆石微量元素特征，以期揭示花岗岩捕虏体的性质，探讨其成因.利用锆石微量元素研究其原岩类型的方法已有报道（Belousova et al.， 2002； Wang et al.， 2012a）.岩浆锆石通常具有明显重稀土富集和Ce正异常，这受控于锆石结晶时的岩浆成分以及锆石的稀土和微量元素分配系数，因而不同岩浆锆石其稀土和微量元素在含量和组成上具有明显区别（赵振华， 2016； 赵志丹等， 2018）.Wang et al.（2012a）通过对喜马拉雅典型S、I、A型花岗岩研究表明，来自S型花岗岩的锆石则表现出较高的Pb浓度和较低的Th/Pb及显著的Eu负异常，而来自I型花岗岩的锆石具有相对较低的Pb浓度和较高的Th/Pb比值，但对于A型花岗岩中的锆石，这些值在I型和S型花岗岩之间过渡.大坪捕虏花岗岩中833 Ma隐伏S型花岗锆石稀土配分曲线具有明显LREE亏损、HREE富集特征，具有明显Ce正异常和Eu负异常，与Wang et al. （2012a）所报道S型花岗岩配分曲线相似（图5a）.根据锆石Th⁃Pb元素二元图解（赵志丹等， 2018），833 Ma捕虏花岗岩锆石同样全部落点于S型花岗岩范围内（图5b），因而可判断黔东大坪捕虏花岗岩属于过铝质S型花岗岩.

4.1.2　大坪捕虏S型花岗岩与梵净山S型花岗岩年龄对比

本文对江南造山带已公开发表的113个新元古代花岗岩年龄数据统计发现，造山带中花岗岩年龄在970~770 Ma，共有970 Ma、905 Ma、 855 Ma、825 Ma和777 Ma五组主要的峰值年龄，其中江南造山带东段岩浆历史最老且最为复杂，西段和中段相对较为集中，以825 Ma的峰值年龄最为集中（Xin et al.， 2017； Deng et al.， 2018； Xia et al.， 2018）及其参考文献）（图6a）.该期花岗岩包括造山带东段休宁（826 Ma）和歙县（838 Ma）岩体（薛怀民等， 2010）；中段约830 Ma的九岭岩基（Sun et al.， 2017； Wang and Wang， 2021）和西段桂北830~815 Ma的三防、元宝山等S型花岗岩岩基（Wang et al.， 2006； Yan et al.， 2021； Yao et al.， 2021）以及黔东北梵净山地区约830 Ma较小的S型花岗岩岩体（Zhao et al.， 2011； 高林志等， 2011； 王敏等， 2011； Xiang et al.， 2020； Lv et al.， 2021）.

本次所研究黔东大坪捕虏S型花岗岩锆石U⁃Pb谐和年龄为（833±2.6） Ma，这与江南造山带内825 Ma的主要花岗岩峰值相一致.大坪捕虏花岗岩与其北东约100 km的梵净山花岗岩二者结晶年龄在误差范围内一致（图6b）（Zhao et al.， 2011； 高林志等， 2011； 王敏等， 2011； Xiang et al.， 2020； Lv et al.， 2021），表明大坪钾镁煌斑岩所捕虏的隐伏S型花岗岩与梵净山地区出露的S型花岗岩可能为同一时期产物.

4.1.3　大坪捕虏S型花岗岩与梵净山S型花岗岩Hf同位素对比

长期以来，研究发现花岗质熔体主要有两种不同来源： （1）幔源基性岩浆的分离结晶； （2）原有地壳的部分熔融（Zheng and Gao， 2021； 郑永飞， 2022）.花岗岩中锆石Lu⁃Hf同位素组成可作为地壳物质来源性质的记录，通常古老地壳部分熔融的花岗岩ε_Hf（t）值为负值，而新生地壳部分熔融的花岗岩ε_Hf（t）值为正值（Wu et al.， 2006b； Zheng et al.， 2007； Tang et al.， 2020）.此外，花岗岩中锆石的亏损地幔模式年龄（T_DM）也可被用于确定岩浆岩的源区，如果锆石Hf的亏损地幔模式年龄（T_DM）值大于锆石结晶年龄，那么可解释为这些花岗岩来自更古老的源区（Vervoort， 2014）.本文大坪钾镁煌斑岩所捕掳的833 Ma隐伏S型花岗锆石 ε_Hf（t）值为-11.4~-2.3，亏损地幔模式年龄（T_DM）在2.5~1.9 Ga范围内.通常造山带中S型花岗岩被认为与变质沉积岩（Collins and Richards， 2008； Zheng and Gao， 2021）部分熔融有关，因而表明大坪捕虏S型花岗岩可能为古老沉积地壳的部分熔融而来.

造山带中岩浆锆石Hf同位素填图可揭示地壳深部物质组成以及古老地壳的空间展布和时空演化（侯增谦和王涛， 2018； Zhang et al.， 2023b）.笔者通过对江南造山带中已公开发表的花岗岩锆石Hf同位素进行二维核密度估算投图研究发现，江南造山带东段、中段和西段三者在Hf同位素组成上存在较大的差异性（图7a）.在江南造山带东段和中段花岗岩锆石ε_Hf（t）值普遍大于0，具有较为年轻的亏损地幔模式年龄.而在江南造山带西段则花岗岩锆石ε_Hf（t）值绝大多数都小于0，具有较为古老的亏损地幔模式年龄（图7a）.江南造山带内花岗岩锆石 ε_Hf（t）自东向西呈明显的减小趋势（图7a），这种变化可能是由于基底物源和花岗质岩熔体来源不同引起，意味着江南造山带东段花岗岩主要由新生地壳部分熔融形成，而西段花岗岩则是由古老地壳部分熔融而来（Wang et al.， 2014）.笔者单独对本次研究的大坪捕虏花岗岩和梵净山已发表花岗岩锆石Hf同位素进行研究表明，二者Hf同位素组成在二维核密度估算图上与江南造山带西段范围大面积重合，且二者Hf同位素组成基本相同（图7b），都为古老地壳部分熔融所形成，而与江南造山带中段由新生地壳部分熔融形成的九岭岩基明显不同.因而，笔者认为大坪捕虏花岗岩与梵净山花岗岩在形成时代、源岩性质上均具有相似性，二者可能在深部共同构成直径长达100 km的花岗岩基.

4.2　花岗岩岩基对江南造山带西段北侧边界的启示

利用锆石微量元素解释其母岩构造背景和物质来源已被广泛用（Belousova et al.， 2002； Yang et al.， 2012； Drabon et al.， 2022； Chen et al.， 2023）.例如，Chen et al. （2023）通过机器学习发现U、Er和Tm是区分在洋、陆地壳环境中结晶锆石的三个重要元素，这可能是由于大陆地壳和弧岩浆中U和Th的富集和Er、Tm等重稀土的亏损是受俯冲沉积物的加入和/或俯冲地壳的重熔所控制（Grimes et al.， 2007），或更厚地壳和高压熔融（Carley et al.， 2022）所造成.此外，在岩浆系统中U和Nb比值受到洋壳熔融深度、变质脱水、幔源组成变化和地壳同化等多因素影响（Grimes et al.， 2015； Drabon et al.， 2022）.当锆石中U/Nb比值小于20时，通常将其定义为典型地幔岩浆锆石，通过Nb相对亏损和U的富集，可以将俯冲环境与幔源环境中的锆石区分开来（Drabon et al.， 2022）.另一方面，锆石中Th⁃U⁃Nb⁃Hf等元素地球化学行为也为判别锆石构造背景提供了可能的手段（Yang et al.， 2012）.相较于板内环境，弧岩浆中Nb的含量明显较少（Sun and McDonough， 1989； Pearce and Peate， 1995），表明在岩浆分馏程度相当的情况下，弧源锆石具有较低的Nb/Hf和较高的Th/Nb值，锆石的地球化学特征为区分造山和非造山岩系提供了重要手段（Yang et al.， 2012）.基于上述理论，可利用锆石中所含微量元素判别该锆石结晶环境进而说明其形成时的构造背景，也可用其作为对火成岩中捕虏物质为“探针”对来探寻地壳深部物质信息.本文通过对大坪捕虏花岗岩锆石log₁₀（U/Tm） vs. log₁₀（Er）研究发现这些锆石基本都形成于大陆地壳环境（图8a）.对锆石U/Nb比值投图显示该花岗岩锆石为明显的弧衍生岩浆（图8b）.大坪捕虏花岗岩锆石中Th⁃U⁃Nb⁃Hf等元素地球化学研究投图表明该花岗岩都与造山作用相关（图8c~8d）.

Yao et al.（2016）对江南造山带东段石角地区江山‒绍兴断裂带内880~850 Ma的石英二长岩、镁铁‒超镁铁质岩、角闪片岩研究显示该断裂带内物质与扬子地块内双溪坞弧火山岩一致，区别于断裂带南部华夏地块的陈蔡杂岩.同时，Yu et al.（2016）通过对中部萍乡张家坊地区早古生代花岗闪长岩研究认为该岩体为华夏地块基底部分熔融形成的I型花岗岩，表明华夏地块与江南造山带地体边界应穿过张家坊岩体附近的西部地区.上述二者通过岩石地球化学证据共同表明江南造山带与华夏地块边界应为江山‒绍兴‒萍乡断裂.而江南造山带北侧与扬子地块边界则有大量地球物理等方法对其进行限定，主要以大庸（张家界）‒九江‒石台断裂来划分（He et al.， 2013； Dong et al.， 2015； Guo and Gao， 2018； 陈昌昕等， 2022； 严加永等， 2022； Han et al.， 2023）.然而，江南造山带西段由于新元古代晚期以来的沉积地层覆盖，仅能通过地球物理方法间接厘定为河池‒张家界一线（Dong et al.， 2015； Guo and Gao， 2018； 严加永等， 2022； 李雪垒等， 2023），缺乏相关蛇绿岩套、弧火山岩以及深部物质等地球化学研究支持.但是河池‒张家界一线并未将梵净山与大坪地区包括在内，桂北和黔东北新元古代褶皱基底（梵净山群、四堡群）和沉积盖层（板溪群、下江群）碎屑物质研究显示二者无论是在物源还是构造背景都相一致（Zhang et al.， 2019），同时梵净山群回香坪组和四堡群文通组玄武岩都同属于同一时期具弧岩浆性质的钙碱性玄武岩（Zhang et al.， 2020），因而将河池‒张家界一线作为江南造山带西段与扬子地块边界并不合适.本文研究的黔东大坪833 Ma捕虏花岗岩表明该花岗岩为与造山作用相关的S型花岗岩，该岩体后期被晚奥陶世（449 Ma， Zhang et al.， 2023a）钾镁煌斑岩带至浅表.该捕虏花岗岩的发现证实了王亮等（2020）基于区域重力所提出的在该区存在隐伏花岗岩的推测，同时说明了黔东地区存在直径超 100 km、成因与造山作用密切相关的隐伏新元古代花岗岩基.根据以上认识，本文认为河池‒张家界一线作为江南造山带西段北侧与扬子的地块边界并不合适，二者界线应以贵阳‒张家界断裂分界.

5 结论

（1）黔东晚奥陶世（449 Ma）大坪钾镁煌斑岩在上升过程中在深部捕虏了新元古代（833 Ma）隐伏S型花岗岩，该花岗岩在锆石年龄和Hf同位素组成上与其北东约100 km的梵净山新元古代花岗岩相似，据此笔者认为在江南造山带西段北侧存在隐伏的新元古代超大花岗岩基.

（2）锆石微量元素研究表明捕虏花岗岩与造山作用相关，通过对比前人对于江南造山带西段与扬子地块边界研究表明，笔者认为江南造山带西段与扬子地块边界为区域上的张家界‒贵阳断裂.

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基金资助

国家自然科学基金项目(42363006)

国家自然科学基金项目(41963006)

国家自然科学基金项目(41603039)