Search
E-mail
RSS

Online submission Editor-in-chief Peer review Staff Login

Journal of Tropical Oceanography >

2023 , Vol. 42 >Issue 1: 124 - 134

DOI: https://doi.org/10.11978/2022021

Anisotropic structure in the back arc region, Taranaki, New Zealand*

  • CAO Lingmin , 1, 2 ,
  • ZHAO Liang 3, 4 ,
  • ZHAO Minghui 1, 2, 4 ,
  • QIU Xuelin 1, 2, 4 ,
  • YUAN Huaiyu 5, 6
Expand
  • 1. Key Laboratory of Ocean and Marginal Sea Geology, South China Sea Institute of Oceanology, Chinese Academy of Sciences, Guangzhou 510301, China
  • 2. Southern Marine Science and Engineering Guangdong Laboratory (Guangzhou), Guangzhou 511458, China
  • 3. State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China
  • 4. University of Chinese Academy of Sciences, Beijing 100049, China
  • 5. ARC Centre of Excellence for Core to Crust Fluid Systems, Department of Earth and Environmental Sciences, Macquarie University, North Ryde NSW 2109, Australia
  • 6. Centre for Exploration Targeting, University of Western Australia, Crawley Perth WA 6009, Australia
CAO Lingmin. email:

Copy editor: YAO Yantao

Received date: 2022-01-31

  Revised date: 2022-04-12

  Online published: 2022-04-08

Supported by

Strategic Priority Research Program of the Chinese Academy of Sciences(XDB42020103)

National Natural Science Foundation of China(42076068)

National Natural Science Foundation of China(91858212)

National Natural Science Foundation of China(91958212)

Key Special Project for Introduced Talents Team of Southern Marine Science and Engineering Guangdong Laboratory (Guangzhou)(GML2019ZD0204)

Fold

Abstract

The Pacific plate is subducting beneath the North Island of New Zealand along the Hikurangi Trench to depths of ~300 km, and is colliding with the South Island to the south, causing clockwise rotation of the North Island. Study on deformation of the mantle wedge in the subduction zone is essential to understand the mantle material movement and its dynamic mechanism. In this study, we investigate the anisotropy in the mantle wedge beneath the Taranaki region in the western backarc area of the North Island using the S-wave splitting measurement of local events with depths ranging from 70 km to 150 km. The results show spatial variations in the fast wave direction and delay time. The NE-SW trending dominant fast direction from the events with depths above 120 km is approximately trench-parallel, reflecting the crystallographic preferred orientation of olivine caused by the trench-parallel mantle flow in the mantle wedge. The events below 120 km depth are mainly from north of the Taranaki region. The predominant fast direction of these events is NNE-SSW, which delay times increase with depth. The Pacific slab steepens abruptly to a near-vertical plane at about 100~150 km depth, which could induce stronger shear deformation of upper mantle material in the deep mantle wedge. Therefore, the NNE-SSW trending anisotropy with larger delay times in the deep mantle wedge north of the Taranaki region may be caused by the combination of trench-parallel mantle flow and strong deformation of deep mantle wedge due to steepening of the dipping Pacific slab. The stronger extension in the deep mantle wedge of the northern backarc is the main reason for the spatial variation of anisotropy.

Key words： Hikurangi subduction zone; Taranaki region; shear wave splitting; seismic anisotropy

Cite this article

CAO Lingmin , ZHAO Liang , ZHAO Minghui , QIU Xuelin , YUAN Huaiyu . Anisotropic structure in the back arc region, Taranaki, New Zealand*[J]. Journal of Tropical Oceanography, 2023 , 42(1) : 124 -134 . DOI: 10.11978/2022021

俯冲带地幔楔变形受到不同的俯冲条件和机制控制, 如板片俯冲的倾角、俯冲板片相对于上覆板块前进或后撤、俯冲板片与上覆地幔或下伏地幔的耦合程度等(Heuret et al, 2005)。了解俯冲带弧后应力状态与地壳上地幔构造及变形关系, 对理解俯冲过程中地幔物质运动的动力学机制至关重要(Uyeda et al, 1979; 宋晓晓 等, 2016)。
新西兰北岛位于太平洋板块与澳大利亚板块的汇聚边界南段之希库兰吉(Hikurangi)俯冲带上(图1), 以北为洋-洋俯冲的汤加—克马德克(Tonga—Kermadec)俯冲带, 以南是一条洋-陆走滑挤压断裂带(Alpine Fault, 位于新西兰南岛)。新西兰北岛西部作为弧后变形类型从北边的弧后伸展到南边以挤压为主的过渡带(Wallace et al, 2004; Nicol et al, 2007) (图1b), 为俯冲带弧后应力变化与地幔楔变形之间的关系研究提供了天然实验场所。以往对新西兰北岛上地幔变形的研究结果表明, 新西兰北岛弧前和岛弧地区的地震各向异性快波方向均平行于海沟走向, 北岛中部的弧后地区表现为弱各向异性或各向同性, 亦或是垂直各向异性, 地幔楔和俯冲板片之下的地幔以平行于海沟方向的运动为主 (Marson-Pidgeon et al, 1999, 2004; Audoine et al, 2004; Greve et al, 2008; Illsley-Kemp et al, 2019; Zal, 2020) (图1c)。然而, 对于弧后南部的塔拉纳基地区的上地幔变形特征的认识仍十分不足, 目前尚未清楚该弧后地区的应力状态和地幔流动是否存在空间分布差异。
图1 新西兰希库兰吉俯冲带构造背景及北岛塔拉纳基地区41台宽频带地震台站分布图
中三角形代表研究中使用的地震台站, 其中蓝色三角形代表最终获得分裂结果的台站; 圆圈代表最终使用的近震震中投影, 不同颜色代表地震的震源深度; TRL(Taranaki-Ruapehu Line)为塔拉纳基—鲁阿佩胡构造线(Stern et al, <a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b43')" rid="b43">1987b</a>)。图b为不同地块边界(粗虚线)的相对运动速率(箭头), 数据引自Wallace等(<a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b51')" rid="b51">2004</a>); 椭圆表示估算值的不确定性; 正、负号分别代表拉张和挤压。图c为研究区及邻区构造背景; 灰色虚线为俯冲板片上界面等深线, 参考Williams等(<a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b54')" rid="b54">2013</a>); 黑色箭头代表板块相对运动(Beavan et al, <a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b5')" rid="b5">2002</a>); 不同颜色条棒代表不同学者计算的远震XKS分裂结果(Marson-Pidgeon et al, <a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b26')" rid="b26">1999</a>; Greve et al, <a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b14')" rid="b14">2008</a>; Zal, <a href="javascript:;" class="mag_content_a" onclick="piaofuRef(this,'b56')" rid="b56">2020</a>); 方框表示研究区范围。该图基于国家测绘地理信息局标准地图服务网站下载的审图号为GS(2016)2937号的标准地图制作

图a Regional tectonic setting of the Hikurangi subduction zone and locations of 41 broadband seismometers used in this study. (a) Triangles represent all the seismometers among which the blue ones yield final reliable splitting measurements. Local events are shown as circles, and colored according to the depth ranging from 70 km to 150 km. The TRL (Taranaki-Ruapehu Line) is marked by the dashed red line. (b) Predicted relative motion (indicated by arrows) across block boundaries (in mm·a-1) made by Wallace and Beavan (2004). Ellipses show the uncertainties of the estimates; plus and minus signify extension and contraction, respectively. (c) Regional tectonic setting of the North Island of New Zealand. Dotted contours of plate interface (in km) are taken from Williams et al (2013). Colored bars represent previous XKS splitting results (Marson-Pidgeon et al, 1999; Greve et al, 2008; Zal, 2020). The black arrow shows the absolute plate motion vector (Beavan et al, 2002). The black rectangle marks the study area

通常, 上地幔变形会引起各向异性矿物的定向排列, 而矿物的定向排列会产生地震波的各向异性(Long et al, 2010)。剪切波分裂方法是获得岩石圈和软流圈地幔介质各向异性特征的重要手段之一。常用的剪切波分裂方法采用的核幔边界转化的剪切波(SKS震相)虽然具备很好的横向分辨率, 但其结果反映的是核幔边界到接收点的积分效应, 对各向异性在垂直深度上的分辨能力较差, 易受多层结构影响 (Silver et al, 1994; Romanowicz et al, 2012)。因此, 以往利用SKS震相获得的新西兰北岛地区NE—SW向快波方向, 在深度上很难分辨地幔楔、俯冲板片内及板下地幔对各向异性结果的贡献。而近震S波分裂, 可约束地震射线采样的地壳或地幔各向异性特征, 不同深度地震的剪切波分裂结果的综合分析可以更有效地解释空间上各向异性的来源与成因。
本文利用塔拉纳基地区台站接收到的深度在70~150km的地方震(来自太平洋板块俯冲引起的俯冲板片界面和界面之下岩石圈地震)的波形数据进行剪切波分裂计算。由于其射线路径采样的空间主要包括地幔楔和上覆板块岩石圈, 这能进一步约束塔拉纳基地区地幔楔的各向异性特征, 为探索希库兰吉俯冲带弧后上地幔的变形特征、成因与动力学机制提供地震学证据。

1 地质构造背景

太平洋板块以小于40~45mm·a-1的速度向新西兰北岛汇聚, 其汇聚方向与希库兰吉海沟走向呈倾斜角度(Beavan et al, 2002; DeMets et al, 2010)(图1)。地质资料、古地磁数据和GPS证据表明, 北岛东部相对于澳大利亚板块经历了顺时针旋转(Wright et al, 1986; Walcott, 1987; Mumme et al, 1989; Wallace et al, 2004; Nicol et al, 2007), 而西部的弧后塔拉纳基盆地的南北部分别表现为伸展和缩短(King et al, 1996; Giba et al, 2010; Reilly et al, 2015)(图1b)。这是由于约25Ma太平洋板块沿希库兰吉海沟向澳大利亚板块下方俯冲, 强烈的俯冲作用使塔拉纳基弧后盆地以挤压缩短为主(Stern et al, 1987a; Stagpoole et al, 2008); 约12Ma 开始受到太平洋板块后撤和与南岛发生碰撞的影响, 新西兰北岛发生顺时针旋转, 旋转速率为~3(°)·Ma-1 (Wallace et al, 2004; Nicol et al, 2007)。同时, 塔拉纳基盆地从北部开始拉张, 并逐渐向南扩张, 以致现今的塔拉纳基盆地表现有北部拉张和南部挤压的构造特征。
塔拉纳基-鲁阿佩胡线(Taranaki-Ruapehu Line, 简称TRL)(Stern et al, 1987b)位于安山岩活火山塔拉纳基火山和鲁阿佩胡火山之间, 呈东西走向, 标志着火山弧的南端终点(图1)。TRL被解释为电导率(Salmon et al, 2011)、地震衰减(Salmon et al, 2011)、重力(Stern et al, 1987a)和地壳厚度(Stern et al, 2013; Dimech et al, 2017) 等地球物理参数的突变带。TRL以北为薄地壳(厚约25km)且低电阻率(<100Ω·m); 以南的结构明显不同, 缺乏火山活动、地壳增厚 (厚约32km)以及上地幔地震衰减(Qp-1)陡变(Stern et al, 2006; Stratford et al, 2006; Eberhart-Phillips et al, 2008; Salmon et al, 2011)。研究认为, 北部地区地幔岩石圈受到强烈拉张发生岩石圈拆沉(Stern et al, 2013; Dimech et al, 2017), 而南部的上地幔较冷(Salmon et al, 2011)。在地震学方面, 贝尼奥夫带的分布显示北岛东部下方俯冲角度较缓, 而俯冲至100~150km俯冲板片变陡至近似垂直(Zal, 2020)。

2 数据与方法

研究中使用了2001年12月至2002年9月期间布设于新西兰北岛弧后塔拉纳基地区的41台临时三分量宽频带地震仪(表1、图1)。这些流动地震台站均采用Guralp CMG-6TD和CMG-40T宽频带地震仪, 采样率为100Hz (Sherburn et al, 2002, 2005)。本文对比了USGS(United States Geological Survey)和GeoNet (http://www.geonet.org.nz)提供的地震目录。由于GeoNet为重定位结果, 且研究中使用的地震的震源深度误差小于4km, 对各向异性结果的影响在可接受范围内, 因此在计算成图分析中采用了GeoNet提供的定位结果(表2)。本研究选取深度70~150km、震级>3级且射线入射角<35°的地震来获得岩石圈下方地幔楔的各向异性特征。入射角<35°是为了避免地表转换波的干扰(Nuttli, 1961)。事件波形数据下载自IRIS (Incorporated Research Institutions for Seismology)数据管理中心。
表1 本研究使用的新西兰北岛塔拉纳基地区41个台站的坐标信息

Tab. 1 List of 41 broadband seismometer stations in the Taranaki region used in this study

台站 东经/° 南纬/° 台站 东经/° 南纬/°
ALBT 174.13 39.15 NOPT 174.49 38.95
AOTT 174.80 39.25 NOWT 174.29 39.61
BRET 174.48 39.32 OMUT 174.02 39.54
BROT 174.19 38.99 OPUT 173.84 39.41
CART 174.02 39.19 PIKT 174.46 39.07
CHET 174.35 39.40 PITT 174.14 39.07
CLMT1 174.24 39.37 PUET 173.91 39.51
CROT 174.24 39.06 PUKT1 174.59 39.04
DOVT 173.91 39.19 PUNT 174.02 39.29
HOLT 174.56 39.43 SAUT 173.94 39.25
INGT 174.46 39.60 STAT 174.35 39.32
INST 174.18 39.54 TAVT 174.46 39.51
KERT 174.25 39.09 TIRT 174.33 39.51
LHUT 174.24 39.45 TOET1 174.36 39.13
LTET 174.61 39.34 TORT 174.45 39.14
LWHT 174.07 39.47 UNGT 173.95 39.33
MAGT 174.61 39.34 WANT 173.97 39.45
MATT1 174.57 39.14 WAST 174.33 39.00
MOKT1 174.58 38.97 WINT 174.13 39.58
MTET1 174.61 39.34 YORT 174.25 39.28
NGAT 174.11 39.24
表2 本研究中最终使用的近震信息

Tab. 2 Information of local events

序号 发震时间 南纬
/°		 东经
/°		 深度/km 深度误差/km 震级
日期 时刻
1 2002-04-14 23:49:43 39.94 174.55 110.7 2.7 3.8
2 2002-04-21 7:23:13 40.08 174.30 92.6 3.9 3.9
3 2002-05-06 20:51:48 39.77 174.59 132.1 3.6 3.8
4 2002-05-16 17:16:12 38.73 175.83 126.2 2.3 5.0
5 2002-05-26 13:05:23 39.15 175.54 99.7 1.9 4.0
6 2002-07-25 14:02:09 39.25 175.02 149.7 2.8 3.9
7 2002-07-27 20:36:20 38.77 175.55 133.4 3.1 4.0
8 2002-08-12 9:10:32 38.60 176.05 132.8 2.4 4.0
9 2002-08-22 17:51:17 39.99 174.11 113.4 3.2 3.5
当剪切波传播经过各向异性介质时会发生分裂, 产生偏振方向垂直、传播速度不同的两个剪切波——快剪切波(快波)和慢剪切波(慢波)(Crampin, 1981)。快波偏振方向与慢波延迟时间是表征各向异性的两个基本参数。本研究采用MFAST(multiple frequency automatic splitting technique)程序(Teanby et al, 2004; Savage et al, 2010; Wessel, 2010), 利用最小特征值法网格搜索到最优的快波偏振方向(φ)和快波与慢波之间的延迟时间(δt)。最小特征值法的优势是可用于波形初始极化方向未知的情况, 利用该算法可以获得初始极化方向(p)及其垂直方向(p⊥)上的波形(Silver et al, 1991)。参考前人在新西兰北岛各向异性的研究(Audoine et al, 2004), 本文采用相同的0.5~3Hz频带对原始记录进行滤波, 以提高S波震相初至的拾取精度, 然后利用MFAST程序完成近震S波分裂计算。通常, 较好的分裂结果具有以下特征: 1) 快波与慢波波形具有很高的相似度; 2) 校正后p⊥分量能量不明显; 3) 校正后质点运动图近似为直线; 4) 计算的最小特征值等值线图收敛。图2为NOPT台站记录的S波分裂结果 的事例。利用该方法最终获得16个地震台站记录的9个地震事件的19个高质量分裂结果(表2、表3和图3)。
图2 台站ΝΟPΤ记录到的近震S波分裂计算事例
为初始极化方向(p)和它的垂直方向(p⊥)的波形, 灰色阴影为选取的时间窗; 图b为校正后p和p⊥的波形; 图c和d分别为校正前、后快波(实线)和慢波(虚线)波形; 图e和f分别为校正前和校正后的质点运动图; 图g为协方差矩阵最小特征值等值线图, 最佳分裂参数为灰色十字标点, 黑色加粗等值线为95%的置信区间。图中右上方显示地震信息与S波分裂结果; 台站NOPT位置参见<xref ref-type="fig" rid="F1">图1</xref>

图a An example of local S-wave splitting measurement using the MFAST program at station NOPT: (a) S phase on incoming polarization direction (p) and its perpendicular value (p⊥) before anisotropic correction; (b) the waveforms after correction; (c) and (d) normalized fast (solid lines) and slow (dashed lines) shear-wave waveforms before and after the splitting correction, respectively; (e) particle motion in the selected window (gray zones in (a)); (f) particle motion in the selected window (gray zones in (b)); (g) results of grid search for optimal φ and δt (cross). Event information and values of splitting measurements are shown at the top right. Location of the station is shown in Fig. 1

表3 近震S波分裂计算结果

Tab. 3 Summary of local S splitting parameters by station

台站 地震序号 Δt/s δt误差/s φ/° φ误差/° 台站 地震序号 Δt/s δt误差/s φ/° φ误差/°
ALBT 4 0.16 0.027 21 6.5 NOWT 2 0.08 0.008 49 8
BROT 4 0.13 0.007 40 2.75 OMUT 2 0.27 0.008 27 6.5
CHET 5 0.12 0.009 73 2.75 OPUT 9 0.10 0.009 70 6.5
DOVT 2 0.11 0.015 77 4 PIKT 5 0.17 0.004 40 1
LWHT 2 0.22 0.003 12 5.5 PITT 6 0.12 0.006 73 5
MOKT1 7 0.14 0.005 4 2 STAT 1 0.11 0.050 31 6.25
MOKT1 8 0.33 0.016 1 1.75 STAT 3 0.06 0.004 17 9.25
NOPT 4 0.20 0.005 5 3.25 TOET1 5 0.12 0.018 39 4.75
NOPT 6 0.19 0.007 14 2.25 TORT 7 0.32 0.003 10 2.25
图3 塔拉纳基地区近震S波分裂结果

图中三角形代表研究中使用的地震台站, 其中蓝色三角形代表最终获得分裂结果的台站。圆代表分裂结果投影在地震震中的位置, 其不同颜色表示不同震源深度; 橘红色线条的方向和长度分别代表分裂结果的快波方向和延迟时间的大小。灰色虚线为地震到接收台站的连线。左上角的玫瑰图分别为震源深度<120km和>120km的地震S波分裂结果的快波方向统计图; 以15°分区, 区域大小正比于该范围内快波方向的数目(玫瑰图下方数字对应半径长度, 代表该快波方向范围内结果个数)。俯冲板片上界面等深线参考Williams等(2013)。墨绿色线条代表SHmax(最大水平挤压应力)方向(Townend et al, 2012); 红色虚线为TRL(位置见图1); 粉色实线(Line N和Line S)为图5中模型剖面的示意位置

Fig. 3 All the S-phase splitting measurement results in the Taranaki region. Triangles represent all the stations used in this study, in which the blue ones yield final reliable splitting measurements. Each red-orange bar is plotted at the location of the event (circle colored according to the depth) and represents a single measurement. Orientation of each bar is parallel to measured fast polarization; length is scaled according to delay time. Raypaths between events and stations are delineated with dotted grey lines. Equal-area rose diagrams of S-phase polarizations of results from events above and below the depth of 120 km are shown in the top left corner of the map, respectively. Sectors are drawn every 15°, with areas proportional to the number of polarizations in those directions. The number to the bottom right of each circle represents maximum petal length (the number of polarizations in the corresponding direction). Dotted contours of plate interface (in km) are from Williams et al (2013). Dark green bars show SHmax orientations calculated from focal mechanisms (Townend et al, 2012). TRL marked as a red dotted line is labeled in Fig. 1. Pink lines (Line N and Line S) denote location of schematic model profiles in Fig. 5

3 计算结果

图3显示了本研究获得的塔拉纳基地区的所有近震S波分裂结果, 每个计算结果都投影在相应的地震震中位置上。所有结果的平均快波极化方向为32.63°, 平均慢波延迟时间为0.16s。图4显示了延迟时间随震源深度的分布以及数目统计分布。最终分裂结果采用的地震深度范围在90~150km之间, 延迟时间主要集中在0.10~0.20s (图4); 且研究区北部地震的平均延迟时间(0.18s)明显大于南部(0.13s), 延迟时间最大值(0.33s)来自最北西端的地震(图3)。
图4 分裂延迟时间与震源深度的关系(a)及慢波延迟时间统计直方图(b)

图a中黑色方框左右两侧线段的长度代表延迟时间的误差大小, 灰色箭头代表延迟时间随深度增加而增大的趋势

Fig. 4 Plots of depth versus delay time (a) and histogram of magnitudes of delay times (b). In (a), black rectangles with black bars denote the delay times and their errors. Grey arrow indicates that delay times increase with depths

将分裂结果按照地震深度来划分, 从120km以上和120km以下地震对应的分裂结果的快波方向玫瑰统计图(图3)中可以看出, 快波优势方向存在一定的差异。120km以上地震的快波优势方向为NE—SW向, 近似平行于俯冲板片等深线走向; 120km以下地震的快波优势方向为NNE—SSW向, 与俯冲方向及俯冲板片等深线走向均存在一定夹角。从延迟时间与地震深度的关系图(图4a)可以看出, 延迟时间随震源深度的增加而有增大的趋势。

4 分析与讨论

通常, 地壳中的各向异性反映了区域应力环境下形成的微裂隙定向排列的方向 (Crampin, 1978; Weiss et al, 1999); 而上地幔中的各向异性主要来源于软流圈中局部地幔对流、绝对板块运动和区域构造运动引起的岩石圈变形(Hess, 1964; Christensen, 1984; Nicolas et al, 1987; Zhang et al, 1995; 高原 等, 2004; Zhao et al, 2005; 常利军 等, 2006)。目前普遍认为上地幔中矿物(如橄榄岩)的晶格优势取向(crystallographic preferred orientation, CPO)是地幔各向异性的主要成因(Hess, 1964; Christensen, 1984; Zhang et al, 1995)。俯冲带各向异性研究显示, 向下俯冲的大洋板块与上覆地幔耦合可以产生地幔楔对流, 表现为弧后地区各向异性快波方向垂直于海沟走向, 通常为A型或C型橄榄岩引起的各向异性。典型的A型橄榄岩形成于低压贫水条件, 而C型橄榄岩形成于低压富水条件(Jung et al, 2001b; Katayama et al, 2004)。在一些俯冲带地区也存在地幔楔快波方向平行于海沟走向, 例如B型橄榄岩形成于高压条件下, 其快波方向垂直于剪应力方向(Jung et al, 2001a; Karato et al, 2008; Long et al, 2013)。此外, 地幔楔中的绕流和俯冲板片边缘的回流等也能造成各向异性快波方向的差异(Long et al, 2008)。前人利用远震SKS震相获得新西兰北岛北部地区的各向异性快波方向以NE—SW向为主, 认为其可能为平行于海沟方向的地幔流动引起的各向异性或B型橄榄岩引起的平行于海沟的各向异性(Audoine et al, 2004); 北岛东部和南部也表现为平行于海沟的快波方向, 主要是平行于海沟的板下地幔流动引起的(Gledhill et al, 1996; Marson-Pidgeon et al, 1999; Greve et al, 2008)。Greve等(2008)发现TVZ (Taupo Volcanic Zone, 陶波火山区; 图1、图3)及以东的SKS分裂快波方向为NE—SW, TVZ中部延迟时间最大(最大可达4.5s), 而TVZ以西延迟时间最小。研究者认为TVZ下方存在较厚的各向异性地幔楔, 较大的延迟时间可能反映了在平行于海沟方向的地幔流作用下, E型橄榄岩(低压环境下形成, 快波方向平行于地幔流动方向) CPO产生的较强的各向异性。TVZ以西的弧后地区表现为弱各向异性或各向同性, 亦或是垂向各向异性(Greve et al, 2008)。北岛西部下方俯冲板片的俯冲角度急剧变陡使得地幔楔远端的动力学过程更为复杂。在TVZ以南, 横跨TRL的地震台阵(N—S向分布的台阵)的XKS分裂结果显示, 平均延迟时间为1.9s, NE—SW向的快波方向可能为地幔楔、俯冲板内及板下地幔中平行于海沟走向的各向异性的叠加(图1c) (Zal, 2020)。
Illsley-Kemp等(2019)利用壳内地震(<40km) S波分裂计算获得塔拉纳基地区及周缘的地壳各向异性特征, 快波方向空间变化明显: 塔拉纳基地区地壳各向异性的快波方向以NE—SW为主, 近似平行于最大挤压应力的轴向方向(图3) (Townend et al, 2012); TRL东段以南的台站的快波方向为E—W向, 与该地区最大挤压应力方向一致。研究区内局部最大挤压应力从北向南由NE—SW向转为E—W向, 发生了顺时针旋转(图3) (Townend et al, 2012)。
本次研究采用近震S波分裂计算观测到的NE—SW或NNE—SSW向的各向异性特征应为地壳和地幔楔中各向异性的叠加, 结合以往的SKS和地壳S波分裂结果进行综合分析, 可以对地幔楔的各向异性特征进行更好的约束。地壳各向异性快波方向以NE—SW向为主(Illsley-Kemp et al, 2019), 与局部最大水平挤压应力方向一致, 也与本研究采用的俯冲相关的较深地震S波分裂的快波方向较相近。根据图4a中S波分裂延迟时间随深度增加而增大的趋势, 认为地幔楔中存在明显的各向异性层。
研究结果表明, TRL以南深度小于120km的地震S波分裂结果以NE—SW向快波方向为主, 与E—W向最大挤压应力方向明显不同, 而与以往观测到的SKS分裂的快波方向一致, 平行于俯冲板片的等深线走向(Gledhill et al, 1996; Marson-Pidgeon et al, 1999; Zal, 2020) (图3); 深度大于120km的地震S波分裂结果只有一个, 快波方向为17°(NNE—SSW向), 且延迟时间较小(0.06s)。TRL以北, 120km之上的地震对应的平均快波方向为50.67°, 平行于俯冲板片的等深线走向; 而深于120km的地震对应的平均快波方向为20.56°, 呈NNE—SSW向, 与120km之上的地震的快波方向存在~30°夹角。110km深度的俯冲板片界面处的剪应力约为170MPa (Stern et al, 1992), 该剪应力条件下可生成A型、C型和E型橄榄岩, 而B型和D橄榄岩需要更高的应力条件(>350MPa)(Katayama et al, 2004)。E型橄榄岩形成于高含水条件下, C型橄榄岩对含水量的要求高于E型橄榄岩(Katayama et al, 2004)。前人的研究认为, TVZ地区较强的各向异性(SKS分裂延迟时间高达4.5s)可能是平行于海沟方向的地幔流动引起的E型橄榄岩CPO产生的(Greve et al, 2008); TVZ以南(SKS分裂平均延迟时间为1.9s)地幔楔和板下地幔存在平行于海沟走向的地幔流动, 引起A型橄榄岩CPO(Zal, 2020)。Zal(2020)观测的TRL以北的SKS分裂延迟时间大于TRL以南的结果, 认为TRL以北存在更厚的各向异性层, 且主要贡献来自于板下地幔, 然而其观测不能对地幔楔的贡献有清晰的认识。本次研究发现, TRL以北的S波分裂平均延迟时间(0.18s)大于南部(0.13s)。地震的深度对地幔楔的各向异性层进行了约束, TRL以北较大延迟时间对应的地震来自大于120km的深度, 且恰好来自高含水的TVZ边缘。因此, 认为延迟时间较大可能与射线路径经过的地幔楔中含水量较高有关(Greve et al, 2008), 且各向异性层厚度较TRL以南更大(图5a)。
图5 沿Line N与Line S剖面地幔楔各向异性模型示意图

剖面位置见图3。图中弧线为距离测线30km以内的地震的代表性射线路径, 其中实线为深度>120km的地震射线路径, 虚线为深度<120km的地震射线路径。图b中的“?”表示对深部地幔楔变形的认识不清

Fig. 5 Schematic of possible anisotropic models along Line N and Line S. Locations of the profiles are shown in Fig. 3. Solid and dotted lines represent raypaths from part of events with depth >120 km and <120 km (within a 30 km width of each profile), respectively. “?” in (b) indicates that deep mantle wedge deformation is unknown due to small delay time from only one event with depth >120 km to the south of TRL

从俯冲板片的形态可以看出, 板片俯冲到100km左右开始迅速变陡(图1)。俯冲板片角度变陡会使深部地幔楔变形更为强烈(图5)。加之太平洋板块与南岛发生碰撞, 使得新西兰北岛顺时针旋转, 也增强了北部的弧后拉张作用。因此, 北部120km以下的地震对应的NNE—SSW快波方向, 可能反映了深部地幔楔的各向异性除了受到平行于海沟走向的地幔流动引起的橄榄岩CPO影响外, 还受深部俯冲板片后撤、俯冲角度加剧变陡引起的地幔变形的影响(图5a)。
TRL以南的结果主要来自深度<120km的地震, 其快波方向平行于俯冲板片的等深线走向, 反映了俯冲板片与上覆地幔楔解耦, 以及地幔楔中存在平行于海沟走向的地幔流动。而本文中, TRL以南只有一个分裂结果来自深度>120km的地震, 其快波方向与TRL以北深度>120km的地震的快波方向一致, 但由于其延迟时间较小, 对南部深度>120km的地幔楔变形的认识还需要更多的工作来进行约束。从TRL南北两侧上地幔楔中各向异性的差异可以看出, 新西兰北岛弧后TRL以北的拉张地区对应的地幔楔变形更为强烈。

5 结论

本研究通过分析新西兰弧后塔拉纳基地区41台宽频带流动地震台站的近震(深度在70~150 km之间)波形资料, 利用S波分裂方法最终获得16个台站的19个各向异性测量结果。结果显示, 平均快波极化方向为32.63°, 平均慢波延迟时间为0.16s; 深度120km之上和之下的地震的快波优势方向分别为NE—SW和NNE—SSW向。本文认为研究区下方NE—SW快波方向的各向异性是平行于海沟的地幔楔物质流动引起的橄榄岩CPO产生的; 而NNE—SSW向各向异性则来自深度大于120km的地震的射线路径上, 且其延迟时间也相对较大。一方面, 地震来自具有高含水量的TVZ边缘, 地幔楔中地幔物质流动方向平行于海沟走向; 另一方面, 由于俯冲角度变陡使地幔楔强烈变形, 两者共同作用导致>120km地震的S波分裂结果的快波方向呈NNE—SSW向。

References

Publishing order | Descend order by publishing year | Descend order by cited within
[1]
常利军, 王椿镛, 丁志峰, 2006. 云南地区SKS波分裂研究[J]. 地球物理学报, 49(1): 197-204.

CHANG LIJUN, WANG CHUNYONG, DING ZHIFENG, 2006. A study on SKS splitting beneath the Yunnan region[J]. Chinese Journal of Geophysics, 49(1): 197-204. (in Chinese with English abstract)

[2]
高原, 刘希强, 梁维, 等, 2004. 剪切波分裂系统分析方法(SAM)软件系统[J]. 中国地震, 20(1): 101-107.

GAO YUAN, LIU XIQIANG, LIANG WEI, et al, 2004. Systematic analysis method of shear-wave splitting: SAM software system[J]. Earthquake Research in China, 20(1): 101-107. (in Chinese with English abstract)

[3]
宋晓晓, 李春峰, 2016. 西太平洋科学大洋钻探的地球动力学成果[J]. 热带海洋学报, 35(1): 17-30.

DOI

SONG XIAOXIAO, LI CHUNFENG, 2016. Geodynamic results of scientific ocean drilling in the western Pacific[J]. Journal of Tropical Oceanography, 35(1): 17-30. (in Chinese with English abstract)

DOI

[4]
AUDOINE E, SAVAGE M K, GLEDHILL K, 2004. Anisotropic structure under a back arc spreading region, the Taupo Volcanic Zone, New Zealand[J]. Journal of Geophysical Research: Solid Earth, 109(B11): B11305, doi: 10.1029/2003JB002932.

DOI

[5]
BEAVAN J, TREGONING P, BEVIS M, et al, 2002. Motion and rigidity of the Pacific Plate and implications for plate boundary deformation[J]. Journal of Geophysical Research: Solid Earth, 107(B10): 2261, doi: 10.1029/2001JB000282.

DOI

[6]
CHRISTENSEN N I, 1984. The magnitude, symmetry and origin of upper mantle anisotropy based on fabric analyses of ultramafic tectonites[J]. Geophysical Journal International, 76(1): 89-111.

DOI

[7]
CRAMPIN S, 1978. Seismic-wave propagation through a cracked solid: Polarization as a possible dilatancy diagnostic[J]. Geophysical Journal International, 53(3): 467-496.

DOI

[8]
CRAMPIN S, 1981. A review of wave motion in anisotropic and cracked elastic-media[J]. Wave Motion, 3(4): 343-391.

DOI

[9]
DEMETS C, GORDON R G, ARGUS D F, 2010. Geologically current plate motions[J]. Geophysical Journal International, 181(1): 1-80, doi: 10.1111/j.1365-246X.2009.04491.x.

DOI

[10]
DIMECH J L, STERN T, LAMB S, 2017. Mantle earthquakes, crustal structure, and gravitational instability beneath western North Island, New Zealand[J]. Geology, 45(2): 155-158.

DOI

[11]
EBERHART-PHILLIPS D, REYNERS M, CHADWICK M, et al, 2008. Three-dimensional attenuation structure of the Hikurangi subduction zone in the central North Island, New Zealand[J]. Geophysical Journal International, 174(1): 418-434.

DOI

[12]
GIBA M, NICOL A, WALSH J J, 2010. Evolution of faulting and volcanism in a back‐arc basin and its implications for subduction processes[J]. Tectonics, 29(4): TC4020.

[13]
GLEDHILL K, GUBBINS D, 1996. SKS splitting and the seismic anisotropy of the mantle beneath the Hikurangi subduction zone, New Zealand[J]. Physics of the Earth and Planetary Interiors, 95(3-4): 227-236.

DOI

[14]
GREVE S M, SAVAGE M K, HOFMANN S D, 2008. Strong variations in seismic anisotropy across the Hikurangi subduction zone, North Island, New Zealand[J]. Tectonophysics, 462(1-4): 7-21.

DOI

[15]
HESS H H, 1964. Seismic anisotropy of the uppermost mantle under Oceans[J]. Nature, 203(4945): 629-631, doi: 10.1038/203629a0.

DOI

[16]
HEURET A, LALLEMAND S, 2005. Plate motions, slab dynamics and back-arc deformation[J]. Physics of the Earth and Planetary Interiors, 149(1-2): 31-51.

DOI

[17]
ILLSLEY-KEMP F, SAVAGE M K, WILSON C J N, et al, 2019. Mapping stress and structure from subducting slab to magmatic rift: crustal seismic anisotropy of the north Island, New Zealand[J]. Geochemistry, Geophysics, Geosystems, 20(11): 5038-5056, doi: 10.1029/2019GC008529.

DOI

[18]
JUNG H, KARATO S I, 2001a. Effects of water on dynamically recrystallized grain-size of olivine[J]. Journal of Structural Geology, 23(9): 1337-1344.

DOI

[19]
JUNG H, KARATO S I, 2001b. Water-induced fabric transitions in olivine[J]. Science, 293(5534): 1460-1463.

DOI

[20]
KARATO S I, JUNG H, KATAYAMA I, et al, 2008. Geodynamic significance of seismic anisotropy of the upper mantle: New insights from laboratory studies[J]. Annual Review of Earth and Planetary Sciences, 36(1): 59-95.

DOI

[21]
KATAYAMA I, JUNG H, KARATO S I, 2004. New type of olivine fabric from deformation experiments at modest water content and low stress[J]. Geology, 32(12): 1045-1048.

DOI

[22]
KING P R, THRASHER G P, 1996. Cretaceous-Cenozoic geology and petroleum systems of the Taranaki Basin, New Zealand[M]. Lower Hutt: Institute of Geological & Nuclear Sciences.

[23]
LONG M D, SILVER P G, 2008. The subduction zone flow field from seismic anisotropy: a global view[J]. Science, 319(5861): 315-318, doi: 10.1126/science.1150809.

DOI PMID

[24]
LONG M D, BECKER T W, 2010. Mantle dynamics and seismic anisotropy[J]. Earth and Planetary Science Letters, 297(3-4): 341-354.

DOI

[25]
LONG M D, WIRTH E A, 2013. Mantle flow in subduction systems: the mantle wedge flow field and implications for wedge processes[J]. Journal of Geophysical Research: Solid Earth, 118(2): 583-606.

DOI

[26]
MARSON-PIDGEON K, SAVAGE M K, GLEDHILL K, et al, 1999. Seismic anisotropy beneath the lower half of the North Island, New Zealand[J]. Journal of Geophysical Research: Solid Earth, 104(B9): 20277-20286.

[27]
MARSON-PIDGEON K, SAVAGE M K, 2004. Shear-wave splitting variations across an array in the southern North Island, New Zealand[J]. Geophysical Research Letters, 31(21): L21602, doi: 10.1029/2004gl021190.

DOI

[28]
MUMME T C, LAMB S H, WALCOTT R I, 1989. The Raukumara paleomagnetic domain: Constraints on the tectonic rotation of the east coast, North Island, New Zealand, from paleomagnetic data[J]. New Zealand Journal of Geology and Geophysics, 32(3): 317-326.

DOI

[29]
NICOL A, MAZENGARB C, CHANIER F, et al, 2007. Tectonic evolution of the active Hikurangi subduction margin, New Zealand, since the Oligocene[J]. Tectonics, 26(4): TC4002, doi: 10.1029/2006TC002090.

DOI

[30]
NICOLAS A, CHRISTENSEN N I, 1987. Formation of anisotropy in upper mantle peridotites‐A review[M]// FUCHSK, FROIDEVAUXC. Composition, structure and dynamics of the lithosphere-asthenosphere system. Washingtom: American Geophysical Union: 111-123.

[31]
NUTTLI O, 1961. The effect of the earth's surface on the S wave particle motion[J]. Bulletin of the Seismological Society of America, 51(2): 237-246.

DOI

[32]
REILLY C, NICOL A, WALSH J J, et al, 2015. Evolution of faulting and plate boundary deformation in the southern Taranaki Basin, New Zealand[J]. Tectonophysics, 651-652: 1-18.

[33]
ROMANOWICZ B, YUAN HUAIYU, 2012. On the interpretation of SKS splitting measurements in the presence of several layers of anisotropy[J]. Geophysical Journal International, 188(3): 1129-1140.

DOI

[34]
SALMON M L, STERN T A, SAVAGE M K, 2011. A major step in the continental Moho and its geodynamic consequences: the Taranaki-Ruapehu line, New Zealand[J]. Geophysical Journal International, 186(1): 32-44.

DOI

[35]
SAVAGE M K, WESSEL A, TEANBY N A, et al, 2010. Automatic measurement of shear wave splitting and applications to time varying anisotropy at Mount Ruapehu volcano, New Zealand[J]. Journal of Geophysical Research, 115(B12): B12321, doi: 10.1029/2010JB007722.

DOI

[36]
SHERBURN S, ALLEN C J, 2002. Technical report on Taranaki seismograph deployment and initial data processing and archiving[M]. Lower Hutt: Institute of Geological & Nuclear Sciences.

[37]
SHERBURN S, WHITE R S, 2005. Crustal seismicity in Taranaki, New Zealand using accurate hypocentres from a dense network[J]. Geophysical Journal International, 162(2): 494-506.

DOI

[38]
SILVER P G, CHAN W W, 1991. Shear wave splitting and subcontinental mantle deformation[J]. Journal of Geophysical Research: Solid Earth, 96(B10): 16429-16454.

[39]
SILVER P G, SAVAGE M K, 1994. The interpretation of shear-wave splitting parameters in the presence of two anisotropic layers[J]. Geophysical Journal International, 119(3): 949-963.

DOI

[40]
STAGPOOLE V, NICOL A, 2008. Regional structure and kinematic history of a large subduction back thrust: Taranaki Fault, New Zealand[J]. Journal of Geophysical Research: Solid Earth, 113(B1): B01403.

[41]
STERN T, SMITH E G C, DAVEY F J, et al, 1987a. Crustal and upper mantle structure of the northwestern North Island, New Zealand, from seismic refraction data[J]. Geophysical Journal International, 91(3): 913-936.

DOI

[42]
STERN T, HOUSEMAN G, SALMON M, et al, 2013. Instability of a lithospheric step beneath western North Island, New Zealand[J]. Geology, 41(4): 423-426.

DOI

[43]
STERN T A, DAVEY F J, 1987b. A seismic investigation of the crustal and upper mantle structure within the Central Volcanic Region of New Zealand[J]. New Zealand Journal of Geology and Geophysics, 30(3): 217-231.

DOI

[44]
STERN T A, QUINLAN G M, HOLT W E, 1992. Basin formation behind an active subduction zone: three-dimensional flexural modelling of Wanganui Basin, New Zealand[J]. Basin Research, 4(3-4): 197-214.

DOI

[45]
STERN T A, STRATFORD W R, SALMON M L, 2006. Subduction evolution and mantle dynamics at a continental margin: Central North Island, New Zealand[J]. Reviews of Geophysics, 44(4): RG4002.

[46]
STRATFORD W R, STERN T A, 2006. Crust and upper mantle structure of a continental backarc: central North Island, New Zealand[J]. Geophysical Journal International, 166(1): 469-484.

DOI

[47]
TEANBY N A, KENDALL J M, VAN DER BAAN M, 2004. Automation of shear-wave splitting measurements using cluster analysis[J]. Bulletin of the Seismological Society of America, 94(2): 453-463.

DOI

[48]
TOWNEND J, SHERBURN S, ARNOLD R, et al, 2012. Three-dimensional variations in present-day tectonic stress along the Australia-Pacific plate boundary in New Zealand[J]. Earth and Planetary Science Letters, 353-354: 47-59.

[49]
UYEDA S, Kanamori H, 1979. Back-arc opening and the mode of subduction[J]. Journal of Geophysical Research, 84(B3): 1049-1061.

DOI

[50]
WALCOTT R I, 1987. Geodetic strain and the deformational history of the North Island of New Zealand during the late Cainozoic[J]. Philosophical Transactions of the Royal Society A Mathematical, Physical and Engineering Sciences, 321(1557): 163-181.

[51]
WALLACE L M, BEAVAN J, MCCAFFREY R, et al, 2004. Subduction zone coupling and tectonic block rotations in the North Island, New Zealand[J]. Journal of Geophysical Research: Solid Earth, 109(B12): B12406.

[52]
WEISS T, SIEGESMUND S, RABBEL W, et al, 1999. Seismic velocities and anisotropy of the lower continental crust: a review[J]. Pure and Applied Geophysics, 156(1): 97-122.

DOI

[53]
WESSEL A, 2010. Automatic shear wave splitting measurements at Mt. Ruapehu volcano, New Zealand[D]. Wellington: Victoria University of Wellington.

[54]
WILLIAMS C A, EBERHART-PHILLIPS D, BANNISTER S, et al, 2013. Revised interface geometry for the Hikurangi Subduction Zone, New Zealand[J]. Seismological Research Letters, 84(6): 1066-1073.

DOI

[55]
WRIGHT I C, WALCOTT R I, 1986. Large tectonic rotation of part of New Zealand in the last 5 Ma[J]. Earth and Planetary Science Letters, 80(3-4): 348-352.

DOI

[56]
ZAL H J, 2020. Seismic anisotropy and velocity structure in the North Island, New Zealand[D]. Wellington: Victoria University of Wellington.

[57]
ZHANG SHUQING, KARATO S I, 1995. Lattice preferred orientation of olivine aggregates deformed in simple shear[J]. Nature, 375(6534): 774-777.

DOI

[58]
ZHAO LIANG, ZHENG TIANYU, 2005. Using shear wave splitting measurements to investigate the upper mantle anisotropy beneath the North China Craton: distinct variation from east to west[J]. Geophysical Research Letters, 32(10): L10309.

Options
Outlines

/

〈 〉