地幔对流产生的作用在岩石圈底部(即岩石圈和软流圈分界面)的剪切拖曳力是板块构造运动的力源之一(Chapple and Tullis, 1977),是造成岩石圈变形、大尺度应力状态和分布格局(Fu 1983,1990;Steinberger et al., 2001;王建和叶正仁,2005)、影响地壳长期运动状态(黄建平等,2008;祝爱玉等,2019)和克拉通演化(Kaban et al., 2015;Paul et al., 2019)的一个重要因素.因此,岩石圈底部地幔对流剪切拖曳力的估计具有十分重要的意义.然而,对它的精确估计是一个难题.
自20世纪60年代以来,对岩石圈底部地幔对流剪切拖曳力的估计从未间断,主要采用了两类地球动力学方法.第一类是基于重力异常驱动的地幔对流模型.在这类模型中,假定了地球上半部分由两层组成:外层是刚性和静止的;内层是黏性的,且具有足够大的尺寸以使得可以忽略下边界的重力效应和常黏滞系数1021 Pa·s,以及在静止的刚性外层和黏性内层之间的分界面满足水静力学平衡.通过这些假定,极大地简化了模型,便可将重力异常和岩石圈底部的剪切拖曳力联系起来,直接估算出一阶近似的岩石圈底部的剪切拖曳力场(Pick et al., 1988;傅容珊,1990).在估算时,可分为3种情况.第一,采用Runcorn(1964, 1967)建立的根据卫星重力位球谐系数来估算地幔对流导致的岩石圈底部的剪切拖曳力的公式(如,Runcorn,1967;Liu et al., 1976;Liu, 1977, 1978, 1979, 1980a, b,1981,1985;黄培华和傅容珊,1982;Fu,1986;楼海等,2000)进行估算;第二,采用在Runcorn方法的基础上考虑了岩石层厚度变化(如,Eshagh and Tenzer, 2015;Eshagh, 2014, 2015;Tenzer and Eshagh, 2015)的公式进行估算;以及第三,采用改进的可利用地表测量的重力数据的Runcorn公式(如,Pick et al., 1988;熊熊等, 2003, 2010)进行估算.利用这类方法,他们估算了全球、区域(如,太平洋板块、亚洲、澳大利亚、非洲板块)和局部区域(如,中亚、青藏高原、天山造山带、蒙古—贝加尔地区和板内火山)之下的岩石圈底部的剪切拖曳力.拖曳力在碰撞带(如,喜马拉雅碰撞带)、俯冲带(如,西太平洋俯冲带)和构造边界带(如,青藏高原和塔里木盆地之间的阿尔金断裂带)的最大量级可达100 MPa,但是在其他区域,其量级一般小于10 MPa.第二类地幔对流模型中的地幔流体可能由地表运动板块、地幔密度异常、地球内部热异常或者它们的组合来驱动.一些研究者利用全球大地水准面异常约束的地幔热对流模型(Fu,1986;傅容珊,1990)或区域均衡重力异常约束的上地幔小尺度地幔对流模型(傅容珊等, 1994a, b;王景赟等,2000)计算了全球(傅容珊,1990)、华北(Fu et al., 1996)、青藏高原(傅容珊等,1998)和天山造山带(王景赟等,2000)岩石圈底部的地幔对流速度场和剪切拖曳力场.他们(Fu,1986;傅容珊,1990;Fu et al., 1996;傅容珊等, 1994a, b,1998;王景赟等,2000)发现在中国及邻区,岩石圈底部的剪切拖曳力一般小于6.0 MPa.在这类地幔对流的计算中,为了简化计算和建立与大地水准面异常的联系,假定了地幔瑞利数不超过临界值的1.5倍(<105),但实际地球地幔的瑞利数远远大于这个值.对于上地幔,瑞利数约为106,而对于全地幔,其值约为107(傅容珊和黄建华,2001).另外,与大多数基于重力异常的地幔对流模型的方法(第一类方法)一样,这类模型也采用了常地幔黏度和常岩石圈厚度的假设.一些研究者利用基于地震速度结构的地幔对流模型预测了全球(Conrad and Lithgow-Bertelloni, 2006;朱涛,2011;Paul et al., 2019)、欧亚板块(Warners-Ruckstuhl et al., 2012)和北美板块(Eaton and Frederiksen, 2007)岩石圈底部的剪切拖曳力,发现其值一般小于8.0 MPa.在这类(基于地震速度结构的)模型中,地幔密度异常是由地震速度异常通过速度-密度转换系数直接转换而来的,一般采用了随深度或者同时随深度和温度变化的黏度结构.特别地,Conrad和Lithgow-Bertelloni(2006)、Paul等(2019)在研究中不但考虑了地幔温度和黏度的横向变化,而且还考虑了岩石圈厚度、温度和黏度的横向变化.
显然,Conrad和Lithgow-Bertelloni(2006)、Paul等(2019)的模型更加接近地球的真实情况.不过,他们采用了全球地幔对流模型,且在估算地幔密度异常时仅采用了全球尺度的地震速度结构,不利于用来研究诸如中国大陆等的区域尺度的岩石圈底部的剪切拖曳力场和应变率场.为了解决这个问题,我们对模型进行了改进,建立了基于Koulakov(2011)的亚洲区域尺度的上地幔地震速度结构和全球尺度的地震速度结构(Grand,2002;Simmons et al., 2010;Ritsema et al., 2011;Houser et al., 2008)相互耦合的区域地幔对流模型.同时,与他们(Conrad and Lithgow-Bertelloni, 2006;Paul et al., 2019)的模型一样,本文的模型既考虑了岩石圈厚度、温度和黏度的横向变化,也考虑了地幔温度和黏度的横向变化.由于本文所采用的亚洲区域地震模型的横向分辨率明显比全球尺度的高,因此本文的模型更有利于用来估算地幔对流导致的中国大陆岩石圈底部的剪切拖曳力和应变率场,进而有助于深化认识和理解中国大陆构造运动、岩石圈应力分布格局、地壳长期运动、地震孕育和发生等的深部动力学环境和过程.
1 地幔对流模型本文采用了Zhu和Guo(2021)建立的剪切波分裂各向异性约束下的地幔对流模型来计算中国大陆岩石圈底部的剪切拖曳力和应变率.它是区域地幔对流模型,其内部充满了不可压缩的、具有无穷大普朗特数和满足Boussinesq近似的牛顿流体(如,傅容珊和黄建华,2001).模型中的地幔流由无净旋转(NNR)参考框架下描述的地表板块运动和估算自全球和区域地震速度异常的地幔密度异常共同驱动(如,Zhu,2018;朱涛,2018).这些地幔对流模型具有足够大的计算域:在纬度方向上从10°N延伸到70°N、经度方向上从60°E延伸到150°E,并且在深度方向上包含了整个地幔,使得计算模型的侧边界离中国大陆任何部分的距离超过了1000 km.这样可以避免侧边界引起的人工回流的影响(如,Liu and Stegman, 2011;Zhu,2014).模型具有192(纬度)×288(经度)×64(深度)个有限单元网格.中国大陆大致位于网格中心.网格在水平方向的分辨率是固定的,约35 km,但在深度方向上的分辨率是变化的:从地表到400 km深度为20 km、400 km到1000 km为40 km、以及1000 km以深的区域约为64 km.模型的上边界(地表)为NNR-MORVEL56模型(Argus et al., 2011)描述的板块运动速度边界,下边界(核幔边界)为应力自由边界.采用已经经过验证的CitcomS程序(Moresi et al., 1996;Zhong et al., 2000, 2008;Tan et al., 2006)进行计算.模型使用了与Zhu(2018)和朱涛(2018)相同的基本参数(表 1).
模型中的黏度η(r, T)是随深度和温度同时变化的,其无量纲形式(如,Huang and Zhong, 2005;Conrad and Lithgow-Bertelloni, 2006;Korenaga,2009;Zhu,2018;朱涛,2018;Paul et al., 2019)可以表述为
(1) |
其中,η(r)为随深度变化的黏度.T为温度且T = Tr + δT(Tr为径向温度;δT为横向变化的温度异常).E是激活能且E=E*/(RΔT).T0为无量纲地表温度且T0=Ts/ΔT.本文使用的黏度η(r, T)是Zhu和Guo(2021)在中国大陆核相剪切波分裂各向异性约束下所获得的.首先,他们(Zhu and Guo, 2021)把Forte等(2010)在与全球地幔对流相关的观测值(非静力学自由空气重力场、核幔边界的椭球率异常和板块运动的水平散度)和冰期均衡调整数据约束下获得的径向黏度结构作为参考模型.通过在参考模型的软流层黏度前乘以系数β的方法获得系列的软流层黏度结构(注:仅改变了参考模型的软流层黏度),从而获得了一系列的随深度变化的黏度η(r);其次,通过速度-密度转换系数把地震波速度异常转换为地幔密度异常.对于地幔,通过状态方程将地幔密度异常转换为地幔温度异常δT.Tr采用了Steinberger和Calderwood(2006)的结果.对于大陆和海洋岩石圈的温度结构,分别采用了误差函数和半空间冷却模型来进行计算.由此便可以通过式(1)获得一系列的随深度和温度同时变化的黏度结构η(r, T);再次,通过地幔对流模型来预测剪切波分裂各向异性的快波偏振方向.考虑到快波偏振方向会随着深度变化(如,高原等,2010),预测的偏振方向是各观测台站(图 1中蓝色线段的中点)下整个软流层中的平均值;最后,通过对比分析预测的偏振方向与最小切向能量方法(Silver 1988,1991)的测量结果(图 1中蓝色线段),确定了不同地震速度结构下的最佳黏度结构.
模型所采用的地幔密度异常由两部分组成.一部分来自区域P波速度异常(Koulakov,2011).它主要覆盖了模型中欧亚大陆的上地幔,且在中国大陆的横向分辨率明显比全球地震模型的高,因此可以更好地用于研究中国大陆的地球动力学问题.在转换中,速度-密度转换系数dlnρ/dlnvp=0.15;另一部分来自4个全球S波速度模型.它们主要用来获得Koulakov(2011)模型覆盖区域之外的计算域的密度异常.这些地震速度模型分别为Grand(2002)模型的升级版TX2011、GYPSUM(Simmons et al., 2010)、S40RTS(Ritsema et al., 2011)和HMSLS06 (Houser et al., 2008),相应的dlnρ/dlnvs=0.2、0.15、0.2和0.2.通过这种方式,将横向分辨率相对较差的全球速度结构和横向分辨率相对较高的区域速度结构在地幔对流模型中实现了“耦合”.由于热中性浮力的“构造岩石圈”中的化学成分差异是造成大陆岩石圈中的地震速度异常的主要原因(如,Jordan,1975),而除了地表板块运动之外,本文的地幔对流仅仅由热浮力所驱动,所以在将地震速度异常转换为密度异常时,没有考虑大陆岩石圈和克拉通地区的地震速度异常,即在大陆岩石圈和克拉通地区,速度异常被强制设置为零.通过状态方程和Calderwood(2000)提供的随深度变化的热扩散系数将地幔密度异常转换为δT.它与Tr结合便获得了地幔温度结构,然后通过式(1)便获得了随温度变化的黏度结构(如,Zhu,2018;朱涛,2018).
模型还考虑了岩石圈厚度和黏度的横向变化(如,Conrad and Lithgow-Bertelloni, 2006;Conrad et al., 2007;Conrad and Behn, 2010).岩石圈厚度由两部分组成.在中国大陆,采用了地震-热岩石圈厚度(安美建和石耀霖,2006),而在中国大陆之外的区域,采用了根据不同地震层析成像结果推断的岩石圈厚度的平均值(Steinberger and Becker, 2018).在计算岩石圈温度结构时,与前人(如,Conrad and Lithgow-Bertelloni, 2006;Conrad et al., 2007;Conrad and Behn, 2010)所采用的方法一样,假定大陆岩石圈年龄为300 Ma、温度随深度的分布符合误差函数且岩石圈底部温度为1350 ℃(如,Kumar and Gordon, 2009;Jiménez-Díaz et al., 2012; Liu and Hasterok, 2016; Zhu et al., 2020);海洋岩石圈温度满足半空间冷却模型,海洋岩石圈年龄来自Müller等(2008).图 2显示了一个穿过日本大陆和华北克拉通地区的黏度剖面.
在计算中,采用的收敛标准是最后两次迭代获得的速度和压力之间的差异同时小于10-4.
2 结果及分析 2.1 岩石圈底部的黏度不同地震速度模型的岩石圈底部的黏度结构(图 3a,c,e,g)具有非常相似的分布特征:在青藏高原、阿拉善块体、天山造山带、准噶尔盆地、塔里木盆地、鄂尔多斯盆地、川滇地区北部、四川盆地及其周缘和长白山地区的黏度相对较高,一般大于1020 Pa·s,而在川滇地区南部、除四川盆地及其周缘之外的华南、华北和除长白山地区之外的东北的黏度相对较低,一般小于5.62×1019 Pa·s.岩石圈底部的地幔黏度存在明显的横向变化.对于GYPSUM、HMSLS06、S40RTS和TX2011模型,横向黏度变化的量级分别为2.08(最大:1.11×1021 Pa·s;最小:9.22×1018 Pa·s;平均:1.61×1020 Pa·s;图 3a)、1.97(最大:1.04×1021 Pa·s;最小:1.10×1019 Pa·s;平均:1.49×1020 Pa·s;图 3c)、2.10(最大:1.04×1021 Pa·s;最小:8.31×1018 Pa·s;平均:1.48×1020 Pa·s;图 3e)和2.10(最大:1.22×1021 Pa·s;最小:9.67×1018 Pa·s;平均:1.62×1020 Pa·s;图 3g).基于这4个地震速度结构预测的中国大陆岩石圈底部的(算术)平均黏度的最大值为1.10×1021 Pa·s,最小值为9.54×1018 Pa·s,以及平均值为1.16×1020 Pa·s(表 2,图 3i),其横向变化的量级约为2.06.也就是说,中国大陆岩石圈底部的黏度的横向变化范围约为2个量级.不同区域的岩石圈底部的黏度分布范围和平均值见表 2.
基于不同地震速度结构的地球动力学模型预测的岩石圈底部的水平速度场的基本特征非常相似(图 3b,d,f,h).地幔对流速率一般不超过4.5 cm·a-1.较高的地幔对流速率(>2.5 cm·a-1)出现在青藏高原中部、天山造山带、塔里木盆地中部、阿拉善块体中部、川滇地区北部、鄂尔多斯盆地、华南西缘和东缘、华北东南缘和北缘、以及除松辽盆地之外的东北绝大部分区域,而剩余区域的较低,一般小于2.0 cm·a-1,尤其在川滇地区南部、华南的江南造山带和华夏块体的过渡带、以及华北南部一般小于1.0 cm·a-1(图 3b,d,f,h).不同的全球S波速度结构会造成岩石圈底部的对流速率的变化.对于GYPSUM、HMSLS06、S40RTS和TX2011模型,对流速率分别为0.04~4.05 cm·a-1(平均:2.07 cm·a-1;图 3b)、0.04~3.84 cm·a-1(平均:2.02 cm·a-1;图 3d)、0.02~4.11 cm·a-1(平均:2.09 cm·a-1;图 3f)和0.07~4.19 cm·a-1(平均:2.13 cm·a-1;图 3h),相应的空间变化范围分别为2.00、1.98、2.31和1.78个量级.基于这4个地震速度结构预测的中国大陆岩石圈底部的(算术)平均对流速率为0.01~3.99 cm·a-1(平均:2.07 cm·a-1;图 3j;表 2),其空间变化约为2.60个量级.这表明,中国大陆岩石圈底部的对流速率具有明显的横向变化,其空间变化范围可达2个量级以上,但平均速率约为2.0 cm·a-1.不同区域的岩石圈底部的水平对流速率分布范围和平均值见表 2.
岩石圈底部地幔流的基本特征是自西向东流动(图 3中黑色箭头),与地表板块运动的方向一致(图 3中的棕色箭头).但是,除了在青藏高原中部、准噶尔盆地、鄂尔多斯盆地南缘和四川盆地的地幔对流方向与地表板块运动的方向基本重合之外,在其他区域,如天山造山带、塔里木盆地、阿拉善块体、川滇地区、鄂尔多斯盆地、除四川盆地之外的华南、华北和东北,地幔对流方向明显偏离了地表板块运动速度的方向,尤其在川滇地区南部、华南南部和东部、华北和东北,这表明了地幔密度异常驱动的地幔对流对整个地幔对流系统的显著影响.也就是说,仅以地表板块运动驱动的简单软流层流动模型可能在探讨青藏高原中部、准噶尔盆地、鄂尔多斯盆地南缘和四川盆地的软流层各向异性成因时是合理的,但在其他区域,可能就失效了.
2.3 岩石圈底部的剪切拖曳力和应变速率球坐标系(r,θ,φ)下剪切拖曳力与应变率的关系:
(2) |
(3) |
其中,ur,uθ,uφ为对流速率,τrθ和τrφ为剪切拖曳力,
从图 4可以看出,与岩石圈底部的对流速度一样,基于不同地震速度结构的岩石圈底部的剪切拖曳力和应变率场的基本特征也是非常相似的.但是,与自西向东的地幔物质流动不同,它们并不是呈现出自西向东拖曳的简单特征.在准噶尔盆地、天山造山带、阿拉善块体、塔里木盆地和华南大部分区域,岩石圈底部的拖曳力的方向主要为由北西向南东;在青藏高原,西端主要为由西向东或由南西西向北东东,然后呈现出发散状态,接着在中东部转变为由北西向南东方向;在川滇地区,从北部的北西向南东方向转变为南部的由西向东或北北西向南南东方向;在鄂尔多斯盆地,主要为由西向东方向;在华北,从北部的由(近)北向(近)南转变为南部的由北西向南东方向;以及在东北,从北部的由北东向南西转变为南部的由北北东向南南西方向.
在天山、帕米尔造山带与塔里木盆地的过渡带、青藏高原的拉萨地块、华南的江南造山带和东北的长白山造山带,剪切拖曳力较大,一般大于1.5 MPa,而其他区域的较小,一般小于1.0 MPa,尤其在准噶尔盆地、塔里木盆地中部、青藏高原的柴达木盆地西缘、巴颜喀拉块体中部偏西和羌塘块体中部区域、川滇地区、华南的南部沿海地区、鄂尔多斯盆地、华北和东北的大部分地区,剪切拖曳力一般小于0.5 MPa.图 4还表明,中国大陆岩石圈底部的剪切拖曳力最大不超过4.0 MPa.对于GYPSUM、HMSLS06、S40RTS和TX2011模型,岩石圈底部的剪切拖曳力分别为0.004~3.21 MPa(平均:0.71 MPa;图 4a)、0.013~3.96 MPa(平均:0.73 MPa;图 4c)、0.002~3.46 MPa(平均:0.64 MPa;图 4e)和0.005~2.86 MPa(平均:0.60 MPa;图 4g),相应的空间变化范围分别为2.90、2.48、3.24和2.76个量级.基于这4个地震速度结构预测的中国大陆岩石圈底部的(算术)平均剪切拖曳力为0.03~3.37 MPa(平均:0.67 MPa;图 4i),其空间变化约为2.05个量级.这表明,虽然中国大陆岩石圈底部的剪切拖曳力具有明显的横向变化,其空间变化约为2个量级,但平均值不超过1.0 MPa.
剪切拖曳力大并不意味着剪切应变率高.图 4表明,在剪切拖曳力较小(<1.0 MPa)的川滇地区南部、华南东缘和华北东南缘,剪切应变率相对较高,一般大于20/100 Ma,而在剪切拖曳力较大的天山、帕米尔造山带与塔里木盆地的过渡带、青藏高原的拉萨地块、华南的江南造山带和东北的长白山造山带,剪切应变率中等,在(15~20)/100 Ma.值得注意的是,在东北的东北部,剪切拖曳力很小(<0.5 MPa),但是剪切应变率却不低,位于中等剪切的范围.对于整个中国大陆而言,岩石圈底部的剪切应变率不超过45/100 Ma.基于GYPSUM、HMSLS06、S40RTS和TX2011的地球动力学模型预测的岩石圈底部的剪切应变率分别为(0.04~36.21)/100 Ma(平均:8.67/100 Ma;图 4b)、(0.20~41.66)/100 Ma(平均:10.09/100 Ma;图 4d)、(0.06~39.35)/100 Ma(平均:9.04/100 Ma;图 4f)和(0.05~29.85)/100 Ma(平均:7.36/100 Ma;图 4h),相应的空间变化范围分别为2.96、2.32、2.82和2.78个量级.基于这4个地震速度结构预测的中国大陆岩石圈底部的(算术)平均剪切应变率的最大值为34.66/100 Ma,最小值为0.47/100 Ma,以及平均值为8.79/100 Ma(图 4j,表 2),其空间变化范围约为1.87个量级.这表明,中国大陆岩石圈底部的平均剪切应变率不超过10/100 Ma,空间变化范围约为2个量级.
3 讨论利用Runcorn(1964, 1967)及其改进方法获得的岩石圈底部的最大剪切拖曳力可达100 MPa的量级(如,黄培华和傅容珊,1982;Eshagh and Tenzer, 2015;Eshagh,2015).不过它们主要出现在边界带如喜马拉雅碰撞带、阿尔金断裂带、西太平洋俯冲带,而在其他区域,剪切拖曳力一般小于10 MPa(Liu,1978;朱岳清等,1984;Fu,1989;傅容珊,1990;熊熊等, 2003, 2010;Eshagh and Tenzer, 2015;Eshagh,2015;Tenzer and Eshagh, 2015).换句话说,在中国的绝大部分区域,该方法获得岩石圈底部的剪切拖曳力小于10 MPa(图 5).特别地,在华北和中国台湾地区,其值分别为2.0~7.5 MPa(朱岳清等,1984)和0~3.0 MPa(Tenzer and Eshagh, 2015)(图 5).力矩平衡分析给出了大陆岩石圈底部和海洋板块下的剪切拖曳力为0.1~4.0 MPa的估计结果(图 5;Chapple and Tullis, 1977),这与中国台湾地区的(0~3.0 MPa;Tenzer and Eshagh, 2015;图 5)相当.运动学和动力学模型预测的全球、欧亚大陆、北美板块的岩石圈底部的剪切拖曳力在0~8 MPa之间(傅容珊,1990;Conrad and Lithgow-Bertelloni, 2006;Eaton and Frederiksen, 2007;朱涛,2011;Paul et al., 2019;图 5),这与利用重力数据获得的结果相当.本文预测的中国大陆岩石圈底部的拖曳力为0~3.37 MPa,其位于前人结果的范围内,但是最大值明显低于一些研究者(如,傅容珊, 1990, 6 MPa;Conrad and Lithgow-Bertelloni, 2006,8 MPa;朱涛, 2011, 6.6 MPa)利用地球动力学模型预测的结果,而与另一些地球动力学模型预测的结果,如Paul等(2019)的4.0 MPa非常接近.
本文获得的岩石圈底部的剪切拖曳力表明,在准噶尔盆地、天山造山带、塔里木盆地、阿拉善块体、青藏高原、川滇地区、鄂尔多斯盆地、华南、华北和东北地区的岩石圈底部的拖曳力分别为0.21~1.26 MPa(平均:0.50 MPa)、0.49~2.11 MPa(平均:1.12 MPa)、0.11~2.23 MPa(平均:1.06 MPa)、0.18~1.50 MPa(平均:0.70 MPa)、0.05~2.17 MPa(平均:0.79 MPa)、0.09~0.78 MPa(平均:0.33 MPa)、0.03~0.90 MPa(平均:0.33 MPa)、0.09~1.91 MPa(平均:0.84 MPa)、0.05~0.75 MPa(平均:0.34 MPa)和0.04~3.37 MPa(平均:0.40 MPa)(表 2,图 6a).由此可以看出,中国大陆西部岩石圈底部的剪切拖曳力相对较高,几乎都大于中国大陆的平均值0.67 MPa,而东部的较低,明显低于整个中国大陆的平均值(表 2,图 6a).青藏高原的平均剪切拖曳力(0.79 MPa)与利用重力和结构模型预测的(0.92 MPa;Eshagh and Tenzer, 2015)比较接近.
前人在估计岩石圈底部的剪切拖曳力时,几乎都假定地幔黏度为常量或者仅仅随着深度变化(如,Runcorn, 1964, 1967;Liu et al., 1976; Liu, 1977, 1978, 1979, 1980a, b,1981,1985;黄培华和傅容珊,1982;朱岳清等,1984;Fu, 1986, 1989;傅容珊,1990;Fu et al., 1996;傅容珊等, 1994a, b,1998;楼海等,2000;王景赟等,2000;熊熊等, 2003, 2010;Eshagh and Tenzer, 2015;Eshagh, 2014, 2015;Tenzer and Eshagh, 2015;Chapple and Tullis, 1977;Warners-Ruckstuhl et al., 2012;Eaton and Frederiksen, 2007),这意味着剪切拖曳力越大的区域,剪切应变率就越高,反之,就越低(参考式(2)—(3)).Paul等(2019)的研究表明,地幔黏度的横向变化会导致剪切拖曳力和应变率呈现负相关,即剪切拖曳力越大,并不意味着剪切应变率就越高,而可能会越低.本文结果支持了Paul等(2019)的结论.因此,边界带,如喜马拉雅碰撞带、阿尔金断裂带、西太平洋俯冲带的大剪切拖曳力,对于常黏度地幔,会导致高剪切应变率(Liu,1978;朱岳清等,1984;Fu,1989;Eshagh and Tenzer, 2015;Eshagh,2015;Tenzer and Eshagh, 2015),但对于横向黏度变化的地幔,可能会获得差异明显、甚至相反的结果.显然,考虑了地幔黏度横向变化的模型,是更符合实际地球情况的.
本文获得的岩石圈底部的剪切应变率表明,在准噶尔盆地、天山造山带、塔里木盆地、阿拉善块体、青藏高原、川滇地区、鄂尔多斯盆地、华南、华北和东北地区的岩石圈底部的剪切应变率分别为(2.38~12.12)/100 Ma(平均:4.99/100 Ma)、(4.04~15.40)/100 Ma(平均:8.93/100 Ma)、(0.47~16.77)/100 Ma(平均:7.95/100 Ma)、(2.46~15.30)/100 Ma(平均:8.76/100 Ma)、(0.62~12.15)/100 Ma(平均:4.81/100 Ma)、(0.72~32.86)/100 Ma(平均:11.12/100 Ma)、(1.06~9.08)/100 Ma(平均:3.81/100 Ma)、(1.04~34.66)/100 Ma(平均:13.94/100 Ma)、(1.65~28.28)/100 Ma(平均:13.24/100 Ma)和(1.14~19.23)/100 Ma(平均:9.40/100 Ma)(表 2,图 6b).对比分析图 4左侧的岩石圈底部的剪切拖曳力和右侧的剪切应变率以及图 6a和图 6b,可以得知,虽然中国西部的岩石圈底部的剪切拖曳力相对较高,但是剪切应变率却不高(<9.0/100 Ma),尤其在青藏高原,其剪切应变率非常低,约为4.81/100 Ma;而在几个剪切拖曳力很小的区域,如川滇地区(0.33 MPa)、华北(0.34 MPa)和东北(0.40 MPa),剪切应变率却很高,分别为11.12/100 Ma、13.24/100 Ma和9.40/100 Ma(表 2),这与Paul等(2019)获得的结果是一致的.
4 结论根据本文的计算结果和分析可知,地幔对流在中国大陆岩石圈底部产生的剪切拖曳力和应变率具有比对流速度更加复杂的分布格局.剪切拖曳力的大小在前人估计的范围之内,具有约2个量级的空间变化,最大值(~3.5 MPa)明显比前人估计的要低.中国西部的剪切拖曳力相对较大,几乎都大于平均值0.67 MPa,而东部的较低,明显低于平均值.中国大陆岩石圈底部的剪切应变率不超过45/100 Ma,空间变化约2个量级.横向变化的地幔黏度可能会导致剪切拖曳力与剪切应变率呈负相关关系,因此虽然在川滇地区南部、华南东缘和华北东南缘的剪切拖曳力较小,但是剪切应变率却相对较高,而在剪切拖曳力较大的天山、帕米尔造山带与塔里木盆地的过渡带、青藏高原的拉萨地块、华南的江南造山带和东北的长白山造山带,剪切应变率并非很高,仅属于本文认为的中等剪切应变率.
附录A 中国大陆地核折射相剪切波各向异性的贡献者Silver和Chan(1991), Makeyeva等(1992), McNamara等(1994), 郑斯华和高原(1994), Hirn等(1995), 丁志峰和曾融生(1996), Guilbert等(1996), Sandvol等(1997), Wolfe和Vernon(1998), Iidaka和Niu(2001), 刘希强等(2001), Liu等(2008), 罗艳等(2004), Chen等(2005, 2010, 2017), Flesch等(2005), Herquel和Tapponnier(2005), Zhao和Zheng(2005), Huang等(2000, 2007, 2008, 2011, 2015), Chang等(2008, 2009, 2011, 2012, 2014, 2015a, b, 2017), 常利军等(2006, 2010, 2015, 2016, 2008a, b, 2012), Lev等(2006), Li和Chen(2006), Sol等(2007), Wang等(2007, 2008, 2013, 2014), 王琼等(2013), Fu等(2008, 2010), Oreshin等(2008), Bai等(2009), 江丽君等(2010), 胡亚轩等(2011), 苗庆杰等(2011), 张丽芬等(2011), 张洪双等(2013), Zhao等(2010), Li和Niu(2010), 李永华等(2010), Li等(2017), 冯强强等(2012), 冯永革等(2016), Soto等(2012), Shi等(2012, 2015), 吴萍萍等(2012), Wu等(2015a, b, 2019), 强正阳和吴庆举(2015), Cherie等(2016), Yu和Chen(2016), Bhukta等(2018), Kong等(2018), Ju等(2019), Mandal(2019).
致谢 非常感谢三位匿名审稿专家为本稿件的改进和完善提出的建设性意见和建议.
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