安山岩的概念主要源于南美安第斯山脉广泛分布的中性火山岩,其主要形成于大陆边缘弧和岛弧等地区,是板块汇聚边缘典型岩石的代表(Gill, 1981; Tatsumi and Eggins, 1995)。高镁安山岩是一种特殊类型的安山岩,其最经典的产地是日本Sectouchi火山带(Tatsumi, 1982; Tatsumi and Ishizaka, 1982)。它比普通的岛弧安山岩具更高的Mg#值、MgO、Cr、Ni和较低的Al2O3和CaO含量及FeOT/MgO比值(Tatsumi, 1982, 2001; Tatsumi and Ishizaka, 1982)。虽然现今高镁安山岩在地球上出露有限,但高镁安山岩的研究一直是国内外岩石学研究领域的热点(Tatsumi and Ishizaka, 1982; Kelemen, 1995; Hirose, 1997; Rapp et al., 1999; Tatsumi, 2001; Xu et al., 2002; Gao et al., 2004; Martin et al., 2005; Wood and Turner, 2009; Zhang et al., 2012b; 邓晋福等, 2010; 张玉芝等, 2015),它们的成因对揭示地壳的生长、地壳拆沉、洋壳的俯冲及流体/熔体与地幔橄榄岩之间的相互作用等过程具十分重要的意义(唐功建和王强, 2010)。绝大部分高镁安山岩分布于汇聚型板块边缘,而只有少量分布于板块内部(Gao et al., 2004; Rao et al., 2008; Wang et al., 2006, 2009; 黄华等, 2007),因此,高镁安山岩的厘定对区域大地构造演化具重要的指示作用。虽然华南板块东南地区侏罗纪岩浆活动强烈(Zhou et al., 2006; Wang et al., 2013a),但目前在右江盆地鲜有侏罗纪岩浆岩报道,特别是高镁安山岩,前人的研究主要集中于盆地晚古生代基性岩和三叠纪基性-中酸性岩(Fan et al., 2008;广西壮族自治区地质矿产局, 1985; 范蔚茗等, 2004; 覃小锋等, 2011; 胡丽沙等, 2012; 李政林等, 2015; 刘寅等, 2015)。本文报道了右江盆地东缘杨屯地区晚侏罗世高镁安山岩,并对其开展了系统的岩相学、年代学和地球化学研究,以查明其岩石成因及动力学机制,进而揭示右江盆地晚侏罗世的大地构造背景。
2 地质背景和样品描述右江盆地在大地构造位置上处于华南板块西南部,是特提斯和滨太平洋构造域的结合部位(曾允孚等, 1995; 杜远生等, 2013)。在地理位置上,该盆地处于桂西、黔西南及滇东南的交接部位,是华南地区重要的多金属矿床富集地之一,广泛发育Au-As-Hg-Sb等低温热液矿床,以卡林型金矿床最为著名(Hu et al., 2002; Su et al., 2008, 2009a, b, 2012; Hu and Zhou, 2012; Chen et al., 2015;贵州地质矿产局, 1987; 广西壮族自治区地质矿产局, 1985; 云南省地矿局, 1990; 胡瑞忠等, 2007; 刘寅等, 2015)。盆地整体在轮廓上呈“菱形”,其东北边界为紫云-南丹断裂,东南端以凭祥-南宁断裂与十万大山盆地相连,西南侧以红河断裂与印支陆块相隔,西北侧以师宗-弥勒断裂与康滇古陆相连(图 1a)。盆地内缺失中元古界-下寒武统和奥陶系-下泥盆统下部地层,泥盆系、石炭系、二叠系和三叠系地层广泛出露于盆地内部,而侏罗系、上白垩统、第三系和第四系地层零星分布于盆地内部,中-晚寒武世地层仅出露于个别地区(广西壮族自治区地质矿产局, 1985; 贵州地质矿产局, 1987; 云南省地矿局, 1990)。该盆地自早泥盆世晚期开始强烈裂陷,形成了世界上罕见的台地-台间相间排列的盆地格局,该格局主要受北西向和北东向的同沉积断裂控制(陈洪德等, 1994; 秦建华等, 1996; 杨怀宇等, 2010; 杜远生等, 2013)。台地相以孤立的浅水碳酸盐岩台地为特征,而台间相以泥质岩、硅质岩等深水沉积为特征(史晓颖等, 2006; Lehrmann et al., 2007; 杜远生等, 2013)。虽然盆地晚古生代-早中生代的构造演化过程和盆地的属性存在强烈争议,但几乎都认为盆地于晚三叠世被重填而闭合,以三叠纪发育巨厚的复理石沉积为特征(曾允孚等, 1995; 秦建华等, 1996; 杜远生等, 2013; 刘寅等, 2015)。右江盆地中生代岩浆活动强烈,基性岩主要出露在隆林-巴马、那坡-靖西-凭祥、八布和富宁以及越东北的高平和Niu Chua等地区,而中酸性火山岩主要分布于峒中、扶隆、凭祥和崇左等地区(Fan et al., 2008;广西壮族自治区地质矿产局, 1985; 范蔚茗等, 2004; 陈懋弘等, 2009; 罗金海等, 2009; 覃小锋等, 2011; 胡丽沙等, 2012; 黄虎等, 2012; 李政林等, 2015; 刘寅等, 2015)。
本文研究区位于右江盆地东缘的武鸣地区,处于紫云-南丹断裂(F2)西南侧和凭祥-南宁断裂(F3)西北侧(图 1a)。区域上出露的地层主要包括石炭系、二叠系、三叠系、侏罗系和白垩系(图 1b)。样品采自武鸣县杨屯村,呈灰绿色,具斑状结构和气孔状构造,斑晶主要为辉石(5%~15%),基质以针状斜长石微晶为主,含少量辉石(图 2a, b)。
本文用于U-Pb定年的锆石全部分选自新鲜的全岩样品,锆石的分选主要采用人工重砂法进行。首先,在双目显微镜下,从用重砂法挑选出的锆石中挑选出自形程度高、无裂隙和透明干净的颗粒。使用环氧树脂将锆石颗粒制作成样品靶,再用磨抛机打磨样品靶直至锆石露出光滑的平面。然后,在显微镜下对锆石进行反射光和透射光照相; 最后,使用扫描电子显微镜上的阴极发光仪详细检查锆石内部结构,以选择最佳分析点。锆石阴极发光成像(CL)分析是在中国科学院地质与地球物理研究所CAMECA SX100型电子探针上完成的,使用的加速电压为15kV。
本文的锆石U-Pb定年分析在中国科学院广州地球化学研究所同位素地球化学国家重点实验室完成。所用仪器包括美国Thermo Fisher Scientific公司的Neptune Plus型多接收电感耦合等离子体质谱仪(MC-LA-ICP-MS)和Resonetics公司的RESOlution M-50型193nm的激光剥蚀系统,关于仪器详细的基本情况参见Zhang et al. (2014)。实验过程中所用的激光束斑直径为24μm,频率为4Hz,能量密度约为5J/cm2。采样方式为单点剥蚀,以He作为剥蚀物质的载气; 采用高纯度的Ar和He气为工作气体。质谱采用离子计数器接收206Pb、207Pb和208Pb的信号,而法拉第杯接收232Th和238U的信号。积分时间设置为0.131s。208Pb气体背景空白小于100cps,238U背景空白小于0.2mv,202Hg背景空白小于2000cps。每个分析点的气体空白采集时间为28s,激光剥蚀时间为30s,共接收450组数据。采用标准锆石Plešovice (337.13±0.37Ma, Sláma et al., 2008)作为外部标准对同位素比值进行校正,采用标准锆石91500 (1062.4±0.6Ma, Wiedenbeck et al., 1995)作为监控样品。实验过程中每测试5个样品点测试2次Plešovice,每测试10个样品点需额外测试2次91500。实验开始和结束时均需测定2次Plešovice和2次91500。数据处理采用ICP-MS Data Cal 8.0软件(Liu et al., 2010)进行分析信号的选择和漂移校正。采用Isoplot (rev. 2.50)程序(Ludwig, 2009)进行U-Pb年龄计算及谐和图绘制。
3.2 主微量元素和Sr-Nd同位素分析首先将无蚀变、无矿化的新鲜岩石样品破碎成拇指大小的碎块,然后置于稀盐酸溶液(浓度约5%)中浸泡至无气泡产生,以淋滤掉碳酸盐矿物,最后用清水冲洗干净并烘干; 将这些烘干的碎块用玛瑙研钵研磨成粒度细于200目的粉末,用于全岩主、微量元素和Sr-Nd同位素分析。全岩主、微量和Sr-Nd同位素分析均在中国科学院广州地球化学研究所同位素地球化学国家重点实验室完成。全岩主量元素分析采用碱熔玻璃片方法,在Rigaku ZSX 100型荧光光谱仪(XRF)上完成,具体的分析流程参见Li et al. (2005)。样品主量元素含量由36种涵盖硅酸盐样品范围的参考标准物质双变量拟合的工作曲线确定。微量元素则用Perkin-Elmer Sciex Elan 6000型ICP-MS进行分析测试,详细的分析流程参见刘颖等(1996)。绝大多数微量元素的分析结果与推荐值之间的相对偏差小于10%,其中多数元素的相对偏差小于5% (刘颖等, 1996)。Sr-Nd同位素分析是在MicroMass Isoprobe型MC-ICP-MS上进行的,该仪器配有九个法拉第杯、四个粒子计数信道和一个电子倍增器共十四个接收器。详细的样品制备和分析流程参见韦刚健等(2002)和梁细荣等(2003)。
4 分析结果 4.1 LA-ICP-MS锆石U-Pb年代学特征本文对采自广西省武鸣县杨屯安山岩样品(13YK-83A)进行了LA-ICP-MS锆石U-Pb定年分析,分析结果见表 1。分选自该样品的锆石呈浅褐或褐色,半透明或透明。大部分锆石呈短柱状,少量呈长柱状,长60~100μm,长宽比为2:1~4:1。锆石CL图像可见典型韵律环带结构,属于岩浆成因锆石。本文测定了18颗锆石,它们的Th/U比值为0.7~2.8,均大于0.1,表明为岩浆成因锆石(吴元保和郑永飞, 2004)。其中两个分析点的207Pb/235U表观年龄分别为2810Ma和738Ma,该两颗锆石应该是岩浆上升过程中捕获的围岩继承锆石。剩余16个分析点的206Pb/238U表观年龄变化范围为153.0~171.2Ma,加权平均年龄为159.3±2.8Ma (MSWD=0.7,n=16,图 3),该年龄代表了杨屯安山岩的喷发年龄。
杨屯安山岩的岩石地球化学分析结果见表 2。样品的SiO2含量变化范围为53.13%~55.71%(不含挥发分),Al2O3含量为14.44%~16.06%,FeOT含量为7.29%~8.29%,具较高的MgO (6.74%~8.85%)和TiO2 (1.17%~1.32%)而较低的CaO (3.24%~7.71%)含量。在TAS图解中,样品落入粗面安山岩范围内(图 4)。样品同时具高K2O (3.39%~4.77%)和全碱(5.97%~7.05%)含量,高K2O/Na2O (1.31~2.33)比值,落入钾玄质岩石区域(图 5a, b)。Tatsumi and Ishizaka (1982)和Tatsumi(1982, 2001)认为高镁安山岩相对于典型岛弧安山岩具更高的MgO (>5%)和更低的FeOT/MgO ( < 1.5);也有学者提出高镁安山岩的SiO2含量为55%~65%,Mg#值大于30 (Kelemen, 1995)。杨屯安山岩具高MgO (6.74%~8.85%)和Mg#值(63~72),低FeOT/MgO (0.83~1.23),均符合上述高镁安山岩的定义(Tatsumi and Ishizaka, 1982; Tatsumi, 1982, 2001; Kelemen, 1995)。因此,杨屯安山岩具高镁安山岩和钾玄质岩石的地球化学特征。
虽然样品具较高的Sr (455×10-6~721×10-6),但相对于典型的埃达克岩(Defant and Drummond, 1990),它们具较高的Y (20.8×10-6~25.6×10-6)和Yb (2.18×10-6~2.40×10-6),而较低的SiO2 (53.13%~55.71%)和Sr/Y (21.2~32.7) (表 2)。样品具较高的相容元素含量,它们的Cr含量为416×10-6~565×10-6,Ni含量为207×10-6~246×10-6 (表 2)。在稀土元素球粒陨石标准化图解上(图 6b),样品明显富集轻稀土元素而亏损重稀土元素,具较弱的Eu负异常(0.82~0.86)。从微量元素原始地幔标准化蛛网图中可知,样品富集Rb、Ba、Th等大离子亲石元素而亏损高场强元素,具明显的Nb-Ta、Sr和Ti负异常(图 6a)。
本文对杨屯安山岩的两件样品进行了Sr-Nd同位素分析,结果见表 2。测得它们的87Sr/86Sr比值为0.70876~0.70879,143Nd/144Nd比值为0.51236~0.51237。采用样品的喷发年龄(159Ma)计算它们的初始Sr-Nd同位素组成,其(87Sr/86Sr)i值变化范围为0.70738~0.70739,εNd(t)值变化范围为-3.6~-3.4,明显不同于云开地区马山和南渡钾玄质岩的同位素组成(图 7; 陈新跃等, 2013; 段瑞春等, 2013; 劳妙姬等, 2015)。
1:20万上林幅区调报告认为在李驴、仙湖、杨屯和六良等地区出露有燕山期辉绿玢岩,它们呈岩墙状产出于白垩纪地层中(广西壮族自治区地质矿产局, 1970)。但该报告也认为由于区域上白垩纪地层露头极差,缺少化石控制,关于其时代归属是臆测的(广西壮族自治区地质矿产局, 1970)。因此,为了确定该套“辉绿玢岩”的形成时代及其围岩地层时代,本文对杨屯地区的“辉绿玢岩”进行了详细的野外地质考察、地球化学及年代学分析。野外和镜下薄片鉴定均表明杨屯地区该套岩石具明显的斑状结构和气孔状构造,样品含辉石斑晶,它们的SiO2含量变化范围为53.13%~55.71% (>53%),该套岩石因此属于典型的安山岩而不是辉绿玢岩。本文样品的LA-ICP-MS锆石U-Pb定年结果(图 3)表明杨屯安山岩的喷发时代为晚侏罗世而不是白垩纪,说明原先臆测的白垩纪地层时代应该属于晚侏罗世。因此,在杨屯地区出露的该套岩石不是原先认为的“辉绿玢岩”而属于安山岩,形成于晚侏罗世; 而李驴、六良等地的“辉绿玢岩”是否也为晚侏罗世安山岩有待进一步研究。此外,杨屯安山岩的厘定为右江盆地存在侏罗纪岩浆活动记录提供了新的年龄证据。
5.2 岩石成因杨屯安山岩具高镁安山岩典型的地球化学特征,目前高镁安山岩主要有以下五种岩石成因:(1) 岩浆混合作用的产物(Kawabata and Shuto, 2005; Guo et al., 2007; Streck et al., 2007);(2) 含水的地幔橄榄岩直接部分熔融的产物(Tatsumi, 1981; Hirose, 1997; Wood and Turner, 2009); (3) 拆沉下地壳部分熔融产生的熔体与地幔橄榄岩相互反应的产物(Kelemen et al., 1998; Xu et al., 2002; Gao et al., 2004; 黄华等; 2007);(4) 俯冲板片部分熔融产生的熔体与上覆地幔橄榄岩反应的产物(Yogodzinski et al., 1994, 1995; Kelemen, 1995; Rapp et al., 1999; Tatsumi, 2001; Tatsumi and Hanyu, 2003; Wang et al., 2006, 2009; Zhang et al., 2012b); (5) 富集地幔部分熔融的产物(Stern et al., 1989; Stern and Hanson, 1991)。
杨屯高镁安山岩的MgO和Nb/La之间无明显的正相关关系(图 8a),它们的Zr/Nb比值不随Nb/La比值的降低而降低(图 8b)。样品的Nb/La (0.31~0.39)和Ce/Pb (3.2~4.7)比值均小于地壳中的平均值(分别为0.40和6.15; Rudnick and Gao, 2003),具低SiO2 (53.13%~55.71%),高Mg#值(63~72),高MgO (6.74%~8.85%)、Cr (416×10-6~565×10-6)和Ni (207×10-6~246×10-6)含量,它们的Nb/La (0.31~0.39)、La/Sm (5.92~6.42)比值和εNd(t)值(-3.6~-3.4)变化不大。这些特征均说明杨屯高镁安山岩不太可能是岩浆混合的产物。实验岩石学结果表明无水的地幔橄榄岩部分熔融无法形成高镁安山岩(Wood and Turner, 2009),而含水的地幔橄榄岩在较低温度( < 1100℃,1GPa)部分熔融可以直接产生高镁安山质岩浆(Tatsumi, 1981; Hirose, 1997),该高镁安山质岩浆具较高的SiO2 (54.35%~60.26%)和Al2O3 (17.24%~21.70%)含量和较低FeOT (4.04%~4.65%)、MgO (5.77%~6.81%)和CaO(8.53%~9.99%)含量(Hirose, 1997)。而本文样品具较低的SiO2 (53.13%~55.71%)、Al2O3 (14.44%~16.06%)和较高的FeOT (7.29%~8.29%)、MgO (6.74%~8.85%)含量,明显不同于上述实验岩石学的结果(Hirose, 1997)。此外,虽然杨屯安山岩具高MgO、Cr和Ni含量,但它们的(87Sr/86Sr)i变化范围为0.70738~0.70739,εNd(t)值为-3.6~-3.4,具富集的Sr-Nd同位素组成(图 7),说明它们不可能是含水的地幔橄榄岩直接部分熔融的产物。
基性下地壳从麻粒岩相向榴辉岩相转变时会导致其密度增加,由于重力作用下地壳会发生拆沉作用(Kay and Kay, 1991; Jull and Kelemen, 2001; Gao et al., 2004),拆沉下地壳部分熔融产生的熔体与地幔橄榄岩相互反应可以形成高镁安山岩(Kelemen et al., 1998; Gao et al., 2004; 黄华等; 2007)。该模式可以解释高镁安山岩具高MgO、Cr和Ni含量的特征,产生的高镁安山岩一般具较高的SiO2、较低的K2O/Na2O和埃达克质岩的地球化学特征(Xu et al., 2002; Gao et al., 2004),而本文样品具较低的SiO2和高K2O/Na2O (1.31~2.33),且不具典型埃达克质岩石的地球化学特征(Defant and Drummond, 1990)。由榴辉岩相部分熔融的熔体通常具Sr和Eu正异常(Stern and Hanson, 1991),而本文样品具明显的Sr和Eu负异常(图 6a, b)。区域上侏罗纪地层并没有经历过高压-超高压变质作用,也缺乏地壳加厚的直接证据(广西壮族自治区地质矿产局, 1985; 贵州地质矿产局, 1987; 云南省地矿局, 1990)。所以,杨屯高镁安山岩也不可能是拆沉下地壳熔融产生的熔体与地幔橄榄岩相互反应的产物。由俯冲洋壳板片熔体与上覆地幔橄榄岩反应产生的高镁安山岩通常具埃达克岩的地球化学特征(Yogodzinski et al., 1994; Rapp et al., 1999; Tatsumi and Hanyu, 2003; Kamei et al., 2004)和类似MORB的同位素组成及高Pb/Nd比值(Class et al., 2000)。本文样品具较低的Pb/Nd (0.47~0.68)比值和明显的Sr、Eu负异常,且不具典型埃达克质岩的地球化学特征,它们的Sr-Nd同位素组成也明显不同于MORB(图 7)。而相比于由俯冲沉积物熔体与上覆地幔橄榄岩反应产生的高镁安山岩(Tatsumi, 2001; Zhang et al., 2012; 张玉芝等, 2015),杨屯高镁安山岩具较低的SiO2和较高的FeOT。区域上也并没有发现同时期的蛇绿岩或与俯冲相关的岩浆岩记录,且尚无地质证据表明在右江盆地存在约160Ma的俯冲板片。因此,同时期的俯冲板片熔体与地幔橄榄岩相互反应不能形成杨屯高镁安山岩。
杨屯高镁安山岩具高K2O和K2O/Na2O,属于典型的钾玄质岩石,说明它们的地幔源区存在金云母和钾质角闪石等富钾矿物(Foley, 1992; Ionov et al., 1997)。含钾质角闪石的地幔橄榄岩部分熔融所形成的熔体通常具较低Rb/Sr ( < 0.06)和较高Ba/Rb (>20)比值,而与含金云母的地幔橄榄岩相平衡的熔体则具较高的Rb/Sr (>0.1)和较低的Ba/Rb ( < 20)比值(Furman and Graham, 1999; Späth et al., 2001)。样品具较高的Rb/Sr (0.20~0.46)和较低的Ba/Rb (7.17~9.30)比值,说明其源区的富钾矿物主要为金云母(图 9a; Turner et al., 1996; Furman and Graham, 1999; Späth et al., 2001; Williams et al., 2004)。杨屯高镁安山岩强烈富集大离子亲石元素和轻稀土元素(图 6a, b),含尖晶石或者石榴子石的地幔橄榄岩部分熔融均形成轻稀土元素强烈富集的熔体(Mckenzie and O’Nions, 1991 )。重稀土元素在尖晶石中分配系数相对较低,故与尖晶石矿物相平衡的熔体一般具较低的Dy/Yb比值和较高的重稀土元素含量,而与石榴子石矿物相平衡的熔体常常具较低的重稀土元素含量和较高的Dy/Yb比值(Blundy et al., 1998; Duggen et al., 2005)。杨屯安山岩具相对平坦的重稀土元素配分模式(图 6b)以及较低的Dy/Yb比值(2.10~2.30),几乎全部落入尖晶石地幔橄榄岩区域(图 9b; Mckenzie and O’Nions, 1991;Miller et al., 1999)。此外,金云母在岩石圈地幔中普遍存在,它们的形成常常与流体/熔体交代岩石圈地幔作用相关(Turner et al., 1996; Rogers et al., 1998; Jiang et al., 2006)。杨屯高镁安山岩具明显富集的Sr-Nd同位素组成(图 7),具明显的Nb-Ta-Ti负异常(图 6a),说明其源区可能受过熔体/流体的交代作用。因此,杨屯高镁安山岩是含金云母的富集岩石圈地幔部分熔融的产物。
扬子陆块和华夏陆块是华南板块的重要组成部分,二者于新元古代碰撞拼贴而形成统一的华南板块(Zhao and Cawood, 1999, 2012; Zhang et al., 2012b; Zhang and Wang, 2016)。华南板块自显生宙以来先后经历了广西运动、印支运动和燕山运动等三期重要的构造-热事件,并伴随着强烈的岩浆活动,其中以中生代岩浆活动最为强烈(Zhou et al., 2006; Wang et al., 2007, 2013a; Shu et al., 2009; Charvet et al., 2010; Zhang et al., 2012a)。晚侏罗世岩浆岩主要分布于华南内陆地区(Zhou et al., 2006; 孙涛, 2006; Wang et al., 2013a),少量出露于东南沿海(Yan et al., 2010; Sewell et al., 2012; Zhang et al., 2015)。一部分学者认为晚侏罗世的岩浆活动与古太平洋板块的俯冲作用紧密相关(Zhou et al., 2006; Li et al., 2007; Jiang et al., 2009, 2015),而另一部分学者则认为它们形成于与古太平洋板块俯冲无关的陆内环境(Wang et al., 2003, 2008, 2013a; Chen et al., 2008)。上述争议的关键在于晚侏罗世古太平洋板块是否已经俯冲至华南板块之下,华南板块底部的岩石圈地幔是否受到了俯冲板片的改造作用等等。
绝大部分高镁安山岩和钾玄质岩石产出于汇聚型板块边缘(Müller et al., 1992; Yogodzinski et al., 1994, 1995; Tatsumi, 2001; Tatsumi and Hanyu, 2003; Zhang et al., 2012b; Liu et al., 2015b; 张玉芝等, 2015),只有极少数高镁安山岩和钾玄质岩形成于板内环境(Gao et al., 2004; Hoang et al., 2009; Muravyeva et al., 2014; 李献华等, 2000, 2001; 陈新跃等, 2013)。杨屯安山岩具高镁安山岩和钾玄质岩石的地球化学特征,其源区为含金云母的富集岩石圈地幔。区域上也没有发现该时期与俯冲相关的岩浆岩记录(Wang et al., 2013a),且右江地区与古太平洋俯冲带相距甚远,说明晚侏罗世右江地区的岩石圈地幔并未受到古太平洋板块俯冲板片的影响。在右江盆地东部的两广交界地区,出露有一系列中晚侏罗世钾玄质岩石(154~166Ma),它们均属于典型的钾玄质岩,形成于板内伸展环境(Li et al., 2004; 李献华等, 2000, 2001; 陈新跃等, 2013; 许华等, 2014; 劳妙姬等, 2015)。在粤西南地区出露的中晚侏罗世高钾钙碱性花岗岩(159~166Ma)是软流圈地幔物质上涌导致地壳物质部分熔融的产物,形成于陆内伸展环境(Huang et al., 2013)。在南岭西部地区出露有晚侏罗世A型花岗岩(付建明等, 2004a, b; Jiang et al., 2009; Huang et al., 2011; Zhou et al., 2015),暗示该地区处于伸展环境。在湘南和桂东南地区出露有中晚侏罗世碱性玄武岩和正长岩,它们具OIB的地球化学特征,同样形成于陆内伸展环境(Li et al., 2004; 王岳军等, 2004)。在相关构造判别图解中,杨屯安山岩样品全部落入板内玄武岩区域内(图 10a, b)。因此,杨屯安山岩更可能形成于板内伸展环境而与岛弧环境无关。右江盆地东缘杨屯钾玄质高镁安山岩的厘定,说明右江盆地晚侏罗世处于板内伸展背景。结合华南板块东部中晚侏罗世岩浆活动特征,反映了华南板块晚侏罗世岩石圈伸展-减薄作用的过程。
(1) 杨屯地区原先被描述成的“辉绿玢岩”属于安山岩,形成于晚侏罗世(159.3±2.8Ma)而不是白垩纪,该地区原先臆测的白垩纪地层时代应为晚侏罗统,属于侏罗纪地层。
(2) 杨屯火山岩具典型高镁安山岩和钾玄质岩的地球化学特征,是含金云母的富集岩石圈地幔部分熔融的产物。
(3) 杨屯安山岩形成于板内伸展环境,暗示晚侏罗世右江地区处于岩石圈伸展-减薄的环境。
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