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地幔氧逸度与俯冲带深部碳循环*

2015-04-13陶仁彪张立飞刘曦

岩石学报 2015年7期
关键词:碳酸盐金刚石熔体

陶仁彪 张立飞 刘曦

TAO RenBiao,ZHANG LiFei and LIU Xi

造山带与地壳演化教育部重点实验室,北京大学地球与空间科学学院,北京 100871

Key Laboratory of Orogenic Belts and Crustal Evolution,MOE,School of Earth and Space Science,Peking University,Beijing 100871,China

2014-09-14 收稿,2014-12-21 改回.

1 地幔氧逸度意义及研究方法

地幔物质的物理和化学属性是现代固体地球科学研究的重要内容。氧逸度作为特殊的成分变量,通过控制地幔矿物中的变价元素,如:过渡金属元素(Sc、Ti、V、Cr、Mn、Fe、Co、Ni、Cu、Zn)和挥发份元素(C、H、O、N、S)的氧化还原状态,影响地幔的矿物组合,固、液相线,流体交代,元素分配等行为(Arculus,1985)。同时,氧逸度还可以显著影响地幔岩石的电导率、扩散性和变形性等物理传导性质(Ryerson et al.,1989;Dai and Karato,2014)。地幔氧逸度的变化也保存和记录了核-幔分异过程和早期地球形成演化信息(Wade and Wood,2005;Wood et al.,2006,2008)。因此,不论是研究地球早期起源和演化历史,还是研究现代固体地幔物质组成及物理化学属性,地幔氧逸度都是极其重要的研究内容。本文重点关注地幔氧逸度对挥发份元素碳(C)的存在形式和迁移属性的影响。譬如,地幔源区氧逸度制约幔源岩浆作用释放含碳气体的种类及含量(Ballhaus and Frost,1994;Holloway,1998)。

地幔氧逸度研究方法主要分为实验室模拟测量、氧逸度计测量和理论计算。其中实验室模拟测量包括了本征氧逸度测量(Sato and Wright,1966;Arculus et al.,1984)以及模拟氧逸度测量(Christie et al.,1986;Wood et al.,1990)。

对于地幔高压样品,由于高压测量过程中样品仓密封性很难得到保证,使得下地幔温压条件下样品的本征氧逸度测量难以进行;而模拟氧逸度测量则由于无法确定体系中Fe3+/∑Fe 值而使其应用受限。Woodland and O’Neill(1997)提出了一种利用惰性贵金属(如Ir)作为氧化还原传感器模拟测量地幔含铁样品在高温高压条件下氧逸度的方法:在含铁地幔样品中加入贵金属(如Ir)作为氧化还原传感器,利用高温高压装置模拟地幔温度、压力条件。实验平衡后淬火,通过分析实验产物中铁和惰性贵金属(Fe-Ir)合金成分,利用实验标定的活度模型计算出体系中对应金属Fe 的活度,从而得到含铁地幔样品在高温高压条件下的氧逸度。最近一些高温高压实验成功应用此法测量了地幔氧逸度的变化及其与碳酸盐稳定性的关系(Stagno and Frost,2010;Rohrbach and Schmidt,2011;Stagno et al.,2011,2013)。这些研究对测量地幔氧逸度和深部碳循环都有重要意义,这种利用惰性贵金属作为氧化还原传感器在高温高压条件下模拟研究地幔氧逸度的方法值得推广研究。

目前应用最为广泛的直接测量地幔氧逸度的方法是氧逸度计法。氧逸度计测量方法是根据幔源岩石中矿物间的氧化还原平衡反应与氧逸度的依存关系,基于相应矿物成分,活度模型,借助热力学计算地幔氧逸度(Haggerty and Tompkins,1983;Luth et al.,1990)。目前关于地幔氧化还原性质的多数数据都是对幔源岩石利用不同氧逸度计计算所得。

地幔氧逸度的理论计算包括两条途径,途径一是:在封闭体系中,若体系的成分、温度、压力均确定,则体系的氧逸度将有确定的取值。根据此原理,通过分析幔源流体包裹体成分及研究流体体系状态方程,建立的利用流体包裹体计算源区氧逸度的方法(Taylor,1990)。理论计算的另一途径是:综合分析来自地幔实际样品的氧逸度数据,结合整个地球物质组成、性质、状态、运动过程及演化历史的认识,建立定量约束地幔氧逸度的模型,根据热力学计算获取地幔任何部位的氧逸度值(Blundy et al.,1991)。刘丛强等(2001)对地幔氧逸度研究方法进行了详细综述,在此不予赘述。

2 地幔氧逸度

2.1 上地幔

上地幔是唯一可以通过天然样品直接测量和估算氧逸度的地球深部系统。通常情况下,来自大陆岩石圈下的上地幔样品会受到部分熔融和交代作用的影响,所以考虑上地幔氧逸度时应当结合部分熔融产生的地幔玄武岩以及残余地幔橄榄岩综合对比来限定上地幔氧化还原状态。

地幔物质部分熔融产生的玄武岩(MORB、IAB、OIB)可以一定程度上反映上地幔的氧化还原状态。前人通过分析快速冷却的洋脊扩张中心枕状玄武岩(MORB)淬火玻璃边的Fe3+/∑Fe 值,获得全球洋中脊扩张中心玄武岩玻璃具有在12 ±2 % 较窄范围内变化的Fe3+/∑Fe 值,并根据经验校准,认为全球平均洋中脊玄武岩玻璃的氧逸度在ΔFMQ-0.41(±0.43)较窄的范围之内(Bézos and Humler,2005)。据此得到的全球MORB 的氧逸度并没有显示与源区属性,熔融程度具有明显的相关性。基于此,Bézos and Humler (2005)提出“缓冲的地幔熔融过程”来解释熔融过程中Fe3+的相容性行为,也就是说MORB 熔体中的Fe2O3可能是受到深部地幔橄榄岩平均氧逸度的控制。与之对比,根据相同的方法获得的全球岛弧玄武岩(IAB)却具有比洋中脊玄武岩(MORB)更高的氧逸度,一般在ΔFMQ 到ΔFMQ + 1.5 范围内变化(Ballhaus,1993a,b)。前人认为这可能是由于俯冲作用引入的地表氧化物质(如H2O)导致岛弧玄武岩具有更高的氧逸度(Bézos and Humler,2005;Kelley and Cottrell,2009)。也有人认为由于很难确定岛弧地幔中Fe 是固有的还是幔源的,所以岛弧地幔的氧化还原性质是很难根据岛弧玄武岩的Fe3+/∑Fe 含量来间接确定(Mallmann and O’Neill,2009)。在此基础上,Lee et al. (2010)发展了用岛弧玄武岩的Zn/FeT值作为岛弧地幔氧逸度计,并提出岛弧地幔的高氧逸度也许并不是由于俯冲氧化性物质的加入引起的,而是由于岛弧岩浆上涌过程中的分异作用引起的,具体分异过程依然不清楚。

上地幔的氧逸度还可以用尖晶石橄榄岩和石榴石橄榄岩氧逸度计来直接限定。对来自上地幔30 ~50km 的尖晶石橄 榄 岩,可 根 据 平 衡 反 应:6Fe2SiO4(橄榄石)+ O2=3Fe2Si2O6(辉石)+2Fe3O4(尖晶石)以及相应矿物成分和活度模型进行尖晶石橄榄岩的氧逸度计算(O’Neill and Wall,1987;Ballhaus et al.,1991)。前人对大量来自不同构造背景下尖晶石橄榄岩进行氧逸度计算并得到从ΔFMQ-3.4 到ΔFMQ+1.5 较大区间变化的氧逸度,这说明软流圈地幔氧逸度在水平空间上具有不均一性(Ballhaus et al.,1991;Ballhaus,1993b;Foley,2010)。地幔石榴石橄榄岩具有比尖晶石橄榄岩更深的来源(50 ~220km)以及相似的全岩Fe2O3含量。然而石榴橄榄岩中石榴石却表现出随压力(来源深度)增加,其Fe3+/∑Fe 值从2%到14% 较大范围的变化(Luth et al.,1990;Canil and O’Neill,1996)。石榴橄榄岩的氧逸度可以根据氧化还原反应:2Fe3Fe32+Si3O12(石榴石)=4Fe2SiO4(橄榄石)+2FeSiO3(斜方辉石)+O2以及相应的矿物成分和活度模型进行计算(Gudmundsson and Wood,1995)。此氧逸度计已经广泛用于限定来自克拉通地幔石榴石橄榄岩的氧逸度。前人通过此氧逸度计计算了大量来自不同地质背景下石榴橄榄岩的氧逸度,发现其具有从ΔFMQ-4.2 到ΔFMQ 的非常大的变化区间(Frost and McCammon,2008)。对于确定成分的石榴橄榄岩,其氧逸度是随着来源深度的增加而降低的(图1曲线1),直到在大于250km 时和镍沉淀氧逸度曲线(NiPC)相交(图1 曲线2)(Frost and McCammon,2008)。上述通过玄武岩和地幔橄榄岩氧逸度计获得的上地幔氧逸度存在明显差异,这可能是由于玄武岩是地幔岩石不同程度熔融作用的产物,并不能直接反应地幔氧化还原属性。故常用的地幔氧逸度数据都是基于尖晶石橄榄岩和石榴石橄榄岩氧逸度计直接限定所得(Frost and McCammon,2008)。但是,玄武岩氧逸度对上地幔氧化还原属性也具有一定的指示意义。

图1 假设富集地幔Fe3+ /∑Fe 为2% 时,在克拉通地热梯度下计算的石榴橄榄岩相对于IW 缓冲剂氧逸度曲线(据Frost and McCammon,2008)Ni 沉淀曲线(NiPC)氧逸度是根据平衡反应Ni2SiO4(橄榄石)=2Ni(金属)+ SiO2(石英) + O2 计 算 而 来. 线1 代 表 平 衡 反 应2Fe3Fe3+2 Si3O12(石榴石)=4Fe2SiO4(橄榄石)+2FeSiO3(辉石)+ O2 计算的石榴橄榄岩的氧逸度曲线;线2 代表石榴橄榄岩氧逸度曲线与NiPC 相交之后的受控于Ni 沉淀曲线的氧逸度变化曲线;线3 代表线1 向金属稳定域的亚稳定延伸线. EMOG/D 氧逸度缓 冲 曲 线 根 据 反 应 MgSiO3(辉石) + MgCO3(菱镁矿) =Mg2SiO4(橄榄石)+C(石墨/金刚石)+O2 计算Fig.1 The oxygen fugacity calculated for a garnet peridotite assemblage assuming a Fe3+ /∑Fe ratio of 2% shown along a cratonic geotherm relative to the IW buffer (after Frost and McCammon,2008)

2.2 上地幔底部和转换带

由于几乎没有来自上地幔深部和转换带的天然样品用于直接计算氧逸度,所以只有通过理论计算(Ballhaus,1995)和高温高压实验(Frost et al.,2004;Rohrbach et al.,2007,2011)去限定上地幔下部和转换带的氧化还原性质。假设全地幔的Fe2O3含量在深度上变化不大,随着压力的增加,石榴石橄榄岩控制氧逸度的反应:2Fe3Fe3+2Si3O12(石榴石)=4Fe2SiO4(橄榄石)+2FeSiO3(辉石)+O2会使促进更多的Fe3+分配进入石榴石矿物相中,使下地幔底部氧逸度随着深度的增加而持续降低。直到250km 左右,石榴橄榄岩氧逸度曲线和镍沉淀氧逸度曲线(NiPC)相交。这时橄榄石中的NiO 会以金属Ni 通过反应:Ni2SiO4(橄榄石)=2Ni(金属)+SiO2+O2沉淀出来,此时深部地幔氧逸度则主要受控于镍沉淀氧逸度曲线(NiPC)(图1)(O’Neill and Wall,1987;Woodland and Koch,2003)。随着压力继续升高,金属铁也会从橄榄石中沉淀出来,这时深部地幔氧逸度则主要通过金属(Fe,Ni)沉淀反应(Ni,Fe)2SiO4(橄榄石)=2(Ni,Fe)(金属)+SiO2+O2来控制。这与天然金刚石包体中发现的金属铁和FexC 碳化物的岩石学观察是一致的(Stachel et al.,1988;Jacob et al.,2004)。这时地幔氧逸度一般低于FMQ 约5 个对数单位。随着Ni 和Fe 金属持续析出,氧逸度应该沿着NiPC 曲线缓慢降低(图1)。从250km 金属Fe 和Ni 析出一直到440km的地幔过渡带,氧逸度只降低0.5 对数单位左右(图1)(Frost and McCammon,2008)。在地幔转换带中,橄榄石相变成瓦德利石(Wadsleyite)和林伍德石(Ringwoodite)使地幔橄榄石含量降低,驱使石榴橄榄岩氧逸度曲线2Fe3Fe3+2Si3O12=4Fe2SiO4+2FeSiO3+O2向右进行,可能会增加转换带计算的氧逸度。然而,Ballhaus (1995)认为在地幔过渡带中,石榴石、瓦德利石以及林伍德石中都具有很强的Fe3+容纳能力,而Fe3Fe3+2Si3O12的活度会由于部分Fe2O3进入瓦德利石和林伍德石(约2%)而导致计算的转换带氧逸度降低。从上述讨论可知,转换带氧逸度的限定取决于Fe3+在石榴石与瓦德利石、林伍德石中容纳能力的差异,这需要更进一步的实验模拟研究来限定。基于地幔对流理论,Ballhaus (1995)假定转换带与软流圈具有相同的Fe3+含量,并通过热力学计算得出从软流圈到过渡带的氧逸度是持续降低的。也就是说过渡带氧逸度随深度的变化依然受控于Fe,Ni 沉淀氧逸度曲线并随深度持续降低。

2.3 下地幔

下地幔(>660km)矿物组合比较均一,主要为钙钛矿型(Mg,Fe)(Al,Si)O3、钙钛矿型CaSiO3和(Mg,Fe)O 方镁石(Fei et al.,1999)。值得注意的是,下地幔主要矿物:钙钛矿型MgSiO3最近首次在天然冲击陨石中被发现,并命名为布里奇曼石(Bridgmanite)(Tschauner et al.,2014)。下地幔MgSiO3钙钛矿结构直到2700km 会转变为后钙钛矿结构(Murakami et al.,2004)。假设下地幔顶部和转换带具有相似全氧含量,就可以根据与金属铁平衡的钙钛矿(布里奇曼石)中Fe3+/∑Fe 最大值来估算下地幔氧逸度区间。Frost et al. (2004)通过高温高压实验得到下地幔钙钛矿结构可以容纳大量的Fe3+(Fe3+/∑Fe 高达0.6),并促使更多的下地幔FeO 通过歧化反应3FeO =Fe +Fe2O3以金属Fe 沉淀出来(Mao and Bell,1977;Frost et al.,2004;Zhang et al.,2014)。如果下地幔全岩的氧含量和上地幔底部基本相同,那么下地幔应该含有1% 的由FeO 歧化形成的金属Fe(Frost and McCammon,2008)。Frost et al. (2004)利用共生的金属Fe 和方镁石成分,得到下地幔氧逸度低于IW 缓冲反应约1.4 对数单位。直到核幔边界,下地幔金属铁、镍为熔融成液态,并在重力作用下趋向与硅酸盐分离进入地核(Wood et al.,2006)。

3 地幔氧逸度的时空演化

从时间演化尺度来讲,地球形成初期45 亿年以前,整个地幔氧逸度都受控于地幔中饱和的金属铁、镍(接近IW 缓冲剂),原始上地幔氧逸度应当低于FMQ 约4.5 个对数单位。前人研究表明上地幔的氧化还原状态在过去的35 亿年以来就一直保持在FMQ 附近(Canil,2002;Li and Lee,2004)。也就是说现今上地幔相对较高的氧逸度应该是在地球形成之后10 亿年演化过程中被逐步升高的(Frost and McCammon,2008)。早期地幔铁、镍金属饱和体系的低氧逸度对一些重要的地球化学属性具有重要影响,譬如地幔中FeO 和亲铁微量元素的含量,大气中挥发份组分以及有机化学键的合成条件等问题(Righter and Ghiorso,2012)。金伯利岩中地幔金刚石包体中经常包裹有CH4和H2流体包体,而这些金刚石中硫化物铅模式年龄一般要比金伯利岩的年龄要老,这也从一方面证实地幔氧逸度是随着时间逐步氧化的。

从空间尺度来讲,Wood et al. (1990)通过对比不同大地构造背景下橄榄岩包体及玄武岩玻璃所记录的氧逸度后得到了以下地幔氧逸度空间分布规律:1)俯冲板块附近的地幔具有最高的氧逸度;2)大洋地幔具有比大陆地幔有更高的氧逸度;3)地幔非熔融区比熔融区域具有更高的氧逸度,且随着部分熔融程度的增加氧逸度减小;4)同深度的地幔柱比地幔其它部位具有更低的氧逸度。并由此认为上地幔氧逸度从地球形成以来逐步氧化的过程有如下几种可能:1)原始上地幔H2O 被地幔饱和金属还原成H2、CH4释放,使上地幔地幔逐步氧化(O’Neill,1991)。2)随着地球逐步变冷,板块俯冲作用开始,俯冲洋壳带入的H2O 和CO2导致上地幔氧化(O’Neill and Wall,1987;Wood et al.,1990;Kasting et al.,1993)。3)地核形成过程中Fe、Ni 金属沉淀,导致下地幔富氧,在大规模地幔对流过程中,形成上地幔的氧化环境(Frost et al.,2004;Galimov,2005;Wood et al.,2008;Rubie et al.,2011)。然而,最近Righter and Ghiorso (2012)计算了氧化平衡反应2Fe+SiO2+O2=Fe2SiO4绝对氧逸度,考虑到温度和压力对氧逸度的影响,认为地核增生可能导致地球氧逸度从高到低变化。地幔氧逸度随时间和空间的演化规律还需要更多的研究来限定。

4 俯冲带深部碳循环与地幔氧逸度

俯冲带作为联系地表和地球深部系统的重要环节,其对研究深部物质(碳)循环乃至能量交换都有重要意义。俯冲带是推动板块构造的主要动力来源,地震学研究已经证实俯冲洋壳可以进入下地幔深度(Fukao et al.,2009)。地球化学研究亦表明俯冲洋壳可进入深部地幔并通过地幔柱作用循环至地表(Hofmann and White,1982;Sobolev et al.,2000)。巴西金伯利岩筒中发现的具有地表碳同位素特征的金刚石及其包裹的深俯冲洋壳物质高压相,从岩石学角度证明了地表碳可随俯冲洋壳进入下地幔深度(>900km),并通过地幔柱岩浆作用循环返回地表(Walter et al.,2011)。在此循环过程中,低氧逸度的地幔还原环境可将俯冲进入深部地幔的地表碳酸盐还原成金刚石,而深部地幔金刚石也可以在地幔物质减压上涌过程中氧化成碳酸岩或者CO2,之后通过地幔柱岩浆作用或者洋中脊岩浆作用循环带出地表(Rohrbach and Schmidt,2011;Stagno et al.,2013)。由此可见,地幔氧逸度对俯冲带深部碳循环具有非常重要的影响。以下我们将结合洋壳形成、水化和碳酸盐化、俯冲脱水脱碳、深俯冲洋壳熔融以及俯冲洋壳通过岩浆作用带出地表的完整循环过程,探讨俯冲带深部碳循环过程中,由于氧逸度的变化导致含碳相表现出不同的存在形式和迁移能力,从而影响深部碳循环过程。

4.1 洋壳形成与地幔碳抽取

软流圈地幔减压熔融产生玄武岩岩浆在洋中脊喷发,形成新生洋壳。地球物理观察表明洋中脊下的熔体可能产生于300km 左右的深度,但是引起熔融的原因还不是很清楚(Team,1998;Gu et al.,2005)。Dasgupta and Hirschmann(2006)通过高温高压实验确定了无水碳酸盐(菱镁矿)化橄榄岩的固相线,认为洋中脊地幔深部300km 的熔体可能是由于地幔中CO2交代引起的。Stagno and Frost (2010)通过高温高压实验重新确定了地幔中碳酸盐熔体和金刚石(石墨)共存的氧化还原条件,并认为地幔中氧逸度在100 ~150km时,就可以让碳酸盐还原成金刚石/石墨稳定存在(图2a)。也就是说,只有在高于150km 的软流圈地幔中,地幔硅酸盐中的Fe3+才可以和石墨(金刚石)通过氧化还原反应形成富碳酸盐的熔体,而非Dasgupta and Hirschmann (2006)在未全面考虑地幔氧逸度对含碳相的影响的情况下认为的300km。那么地球物理观察到的洋中脊下300km 部分熔融可能需要其它的解释,当然也有可能是并不是所有的地幔碳酸岩在150km 时就可以被还原成石墨,或者金刚石并引发碳酸盐化地幔岩石的熔融。譬如在具有高氧逸度的俯冲板片中的碳酸岩可能在300km 时并未完全还原成金刚石,依然可以交代地幔产生熔体。最近,Stagno et al. (2013)通过高温高压实验提出了更好的确定石榴石橄榄岩氧逸度的平衡反应:2Ca3Fe2Si3O12(石榴石)+ 2Mg3Al2Si3O12(石榴石)+ 4FeSiO3(单斜辉石)=2Ca3Al2Si3O12(石榴石)+4Fe2SiO4(橄榄石)+6MgSiO3(单斜辉石)+O2,并指出洋中脊地幔石榴石橄榄岩氧逸度在绝热减压过程中会持续升高(图2a)。也就是说在软流圈减压上涌形成洋脊玄武岩岩浆的过程中,这些深部地幔(>150km)的石墨、金刚石将被软流圈上部硅酸盐中Fe3+氧化成碳酸岩并被带出地表。此过程也说明在新生洋壳形成过程中,地球深部还原性碳(石墨,金刚石)可通过氧化作用从地幔中以CO2抽取并以氧化性的碳酸岩、CO2在洋中脊喷发出地表,形成新生洋壳(图2b)。

图2 绝热上涌洋中脊地幔中含碳相物相变化(据Stagno et al.,2013 修改)(a)硅酸盐地幔氧逸度估算以及不同含碳相稳定域. 灰色直线代表不同Fe3+ /∑Fe 天然石榴石橄榄岩控制的氧逸度演化曲线. 其中黑线代表石金刚石(石墨)和碳酸盐熔体共生时的氧逸度曲线,低压时该曲线降低是由于硅酸盐熔体的出现(Stagno and Frost,2010). 绿色区域代表MORB 氧逸度区间.蓝色箭头代表在金刚石稳定域中含有4%的Fe3+ /∑Fe 的深部地幔向碳酸盐熔体稳定域减压上涌时Fe3+ /∑Fe变化导致的氧逸度变化. 在150km 处,富含金刚石的深部地幔通过氧化熔融作用(Fe2O3 +C =FeO +CO2),消耗硅酸盐中Fe3+,氧化金刚石并产生碳酸盐熔体,而后碳酸盐熔体随洋中脊减压熔融的硅酸盐熔体带出地表. (b)沿着a 中蓝色演化曲线在洋中脊地幔发生的氧化还原熔体产生过程.氧化还原作用可以先于碳酸盐化橄榄岩固相线(300km)在150km 产生碳酸盐熔体Fig.2 Speciation of carbon in adiabatically upwelling mantle (after Stagno et al.,2013)

4.2 新生洋壳水化和碳酸盐化

新生洋壳在洋中脊形成后被海水水化和碳酸盐化(Tatsumi and Eggins,1995;Maruyama and Okamoto,2007)。在洋壳新生过程中,海水沿着深部构造裂隙贯入与洋壳超基性岩和基性岩发生一系列水-岩反应形成蚀变洋壳。期间大量的海水以含水矿物(蛇纹石、角闪石、绿帘石、绿泥石以及粘土矿物)保存在于蚀变洋壳中(Tatsumi,1989;Tatsumi and Eggins,1995)。在洋壳超基性岩蚀变过程中还会释放还原性气体(H2),而其中的橄榄石、辉石等矿物中的FeO 则被氧化成磁铁矿残留在新生蛇纹岩中,使水化的洋壳超基性岩氧逸度明显增加(Berndt et al.,1996;Sleep et al.,2004;黄瑞芳等,2013)。海水中的溶解的CO2可以与洋壳超基性岩蛇纹岩化过程中产生的H2结合,通过费-托合成反应形成无机甲烷(CH4)以及其它的无机碳氢化合物(McCollom and Seewald,2001;Proskurowski et al.,2008;黄瑞芳等,2013;McCollom,2013)。譬如前人证实土耳其Tekirova 蛇绿岩每年大约释放了150 ~190t 的无机CH4气体(Etiope et al.,2011)。洋壳基性岩和超基性岩通过水岩反应形成的Ca2+、Mg2+、Fe2+等离子也可以和海水中碳酸根离子在蚀变洋壳中沉淀出碳酸盐,使海水中CO2以碳酸盐形式沉淀在蚀变洋壳的玄武岩和超基性岩中(Alt and Teagle,1999;Sleep and Zahnle,2001;Kelley et al.,2005;Slagle and Goldberg,2011;Rosenbauer et al.,2012)。洋壳的碳酸盐化蚀变作用可能是一个潜在的固定海水和大气中CO2的地质碳储库(Kelemen and Matter,2008;Kelemen et al.,2011)。新生洋壳继续扩张,由于陆壳来源的碳酸盐沉淀作用以及部分生物作用,可以在蚀变洋壳上部沉淀出含有大量碳酸盐的沉积物乃至大理岩层,使更多的地表CO2以碳酸盐形式固定在洋壳中(Plank and Langmuir,1998)。因此,新生洋壳的水化和碳酸盐化过程可以使大量的碳以碳酸盐形式沉淀在洋壳沉积岩、玄武岩以及超基性岩中。Dasgupta (2013)通过综合前人岩石学研究估算,认为每年大约有5.4 ×1013~8.8 ×1013g 的地表碳通过碳酸盐化作用进入新生洋壳。在洋壳蚀变过程中形成的大量含水矿物中的结晶水、蛇纹岩中磁铁矿以及沉积碳酸盐都是相对氧化性的物质,若是它们伴随俯冲作用进入深部地幔楔,就可能成为形成地幔楔较高氧逸度的俯冲氧化性物质(O’Neill and Wall,1987;Wood et al.,1990;Kasting et al.,1993)。而这些氧化性碳酸岩伴随俯冲作用进入地幔时,可能不会在150km 处完全还原成金刚石或者石墨,而会随着俯冲洋壳进入更深地幔并交代地幔橄榄岩形成碳酸盐化熔体(Dasgupta and Hirschmann,2006)。

4.3 俯冲脱水和脱碳作用

在洋壳俯冲过程中,蚀变洋壳中的水和碳以含水矿物以及碳酸盐矿物形式伴随俯冲板片进入地幔(Poli et al.,2009;Schmidt and Poli,2013)(图3)。在浅层地壳分布最广的钙质碳酸盐(方解石、文石和白云石)在俯冲进变质过程中往往通过流体溶解作用脱出板片(Frezzotti et al.,2011;Manning et al.,2013)。高温高压实验和深部地球岩石样品中仅存的菱镁矿包体都证明只有菱镁矿-菱铁矿等铁镁碳酸盐矿物可以稳定存在到地幔深度(Irving and Wyllie,1975;Biellmann et al.,1993;Yang et al.,1993;Zhang and Liou,1994;Wang et al.,1996;Fiquet et al.,2002;Isshiki et al.,2004;Brenker et al.,2007;Tao et al.,2013)。前人通过热力学模拟计算和岩石地球化学研究都指出俯冲带脱碳过程主要通过俯冲含水矿物脱水过程中形成的化学性流体溶解洋壳碳酸岩而脱出俯冲板片(Kerrick and Connolly,1998,2001a,b;Gorman et al.,2006;Marín-Cerón et al.,2010)。最近,Ague and Nicolescu (2014)通过碳氧同位素、岩石地球化学、流体包裹体等手段综合分析了Cycladic 俯冲带折返大理岩中流体通道特征,发现钙质碳酸盐在靠近流体通道的位置明显降低,质量平衡计算表明有高达60% ~90%的碳伴随流体作用脱出。俯冲脱水溶解碳酸盐形成的C-O-H 流体随后交代地幔楔,以岛弧岩浆作用喷发返回地表(Sleep and Zahnle,2001;Poli and Schmidt,2002)。部分经历高压超高压变质的深俯冲碳酸盐可以通过构造折返作用带出地表,以碳酸盐化榴辉岩和碳酸盐化泥质变质岩出露于高压超高压变质带中,成为研究深部碳循环的理想天然样品(Zhang and Liou,1994,1996;Omori et al.,1998;Messiga et al.,1999;Zhang et al.,2002,2003a;Zhu and Ogasawara,2002;朱永峰,2005;刘福来等,2006;Zhu et al.,2009;Proyer et al.,2013)。最近一些研究也显示,俯冲带折返退变质过程也可能是个重要脱碳过程,研究退变质作用可以帮助理解退变质流体是否将大量的CO2从深部地幔带入浅层水圈或者大气圈(Hazen and Schiffries,2013)。伴随俯冲作用进入深部地幔的碳酸盐化洋壳,成为地表和地球深部碳循环的重要连接点,同时也是改变地幔氧逸度,地球化学等属性的重要介质(Hofmann and White,1982;Wood et al.,1990;Kelley and Cottrell,2009)。

4.4 俯冲流体交代地幔楔

图3 伴随俯冲洋壳高压变质作用过程的脱挥发份(碳和水)作用示意图(据Poli et al.,2009 修改)Fig.3 Schematic illustration showing the complexities in fluid speciation,from ocean floor metamorphism to volatile release at high pressure (after Poli et al.,2009)

俯冲板片中含水矿物分解产生流体溶解碳酸盐,形成CO-H 流体交代地幔楔,产生岛弧岩浆喷出地表(图3)。俯冲进入地幔楔的碳和水主要以H2O 和CO2的形式伴随岛弧岩浆作用循环出地表,形成俯冲-岛弧系统短周期碳循环过程(Tatsumi and Eggins,1995;Tatsumi,2005)。基于岛弧岩浆中常见的氧化性的CO2-H2O 流体事实以及岛弧岩浆流体主要来自俯冲脱水流体的假设,前人一直以来认为俯冲带C-OH 流体是以H2O 和CO2主要的氧化性流体(Wood et al.,1990;Parkinson and Arculus,1999)。然而,近来在西南天山俯冲带榴辉岩和泥质片岩中发现了普遍存在的石墨(Lü et al.,2008,2009,2013)和一些碳氢化合物流体包体(Tao et al.,2015)都说明俯冲带C-O-H 流体并不一定都是前人所认为的氧化性的。Song et al. (2009)在造山带方辉橄榄岩包体中发现了共生的CH4+H2+C 包裹体,并通过碳和稀有气体同位素证实这些还原性流体可能是俯冲洋壳脱水脱碳形成的流体。这种俯冲脱水脱碳产生的还原性流体,与前人认为的俯冲带氧化性富CO2流体的性质是很不一致的:假如俯冲带具有偏低的氧逸度(如西南天山),可能会将部分碳酸盐还原成石墨/金刚石稳定在俯冲带中,并不会随熔融或者流体作用带出俯冲板片,使较多的碳通过俯冲作用进入深部地幔。如果俯冲带氧逸度低到可以将部分俯冲含碳相还原成CH4等易迁移的还原性流体,则如同氧化的CO2流体一样,其会通过流体作用脱出俯冲板片,交代地幔楔并通过岛弧岩浆作用循环出地表。少数岛弧中发现的CH4等还原性气体也许就是来自俯冲带还原性流体(Fiebig et al.,2004)。在俯冲起始阶段,俯冲带流体也可以交代地幔楔橄榄岩,产生蛇纹岩和大量还原性碳氢化合物。部分地幔石榴橄榄岩也可能被来自俯冲带的低温流体交代形成碳酸盐矿物(Yang et al.,1993)。关于俯冲带流体的氧化还原性质,及其对地幔楔的交代作用都需要进一步深入研究。

4.5 俯冲碳酸盐在深部地幔的氧化还原作用

图4 地幔氧化还原环境引起的碳酸岩(熔体)的氧化还原冻结和熔融作用(据Rohrbach and Schmidt,2011 修改)俯冲碳酸盐(熔体)在250km 金属饱和氧逸度条件下还原成稳定的金刚石固定下来;而深部地幔的金刚石也可以通过减压上涌的氧化作用形成易迁移的碳酸盐熔体循环出地表Fig.4 Carbonatitic redox freezing and redox melting caused by redox capacity changes in Earth’s mantle (after Rohrbach and Schmidt,2011)

当俯冲的碳酸盐化洋壳物质进入深部地幔,其液相线与地幔地热线相交,碳酸盐化榴辉岩和泥质变质岩会发生熔融,形成相应的碳酸岩熔体和硅酸盐熔体,这一问题已被许多实验岩石学研究证实(Dasgupta and Hirschmann,2006;Thomsen and Schmidt, 2008; Luth, 2009; Keshav and Gudfinnsson,2010;Litasov and Ohtani,2010;Grassi and Schmidt,2011;Tsuno and Dasgupta,2011)。Dasgupta et al.(2013)通过对比碳酸盐化橄榄岩和无水橄榄岩的实验熔融曲线,认为全球尺度上,碳酸盐化的橄榄岩熔体产生应当在250 ~200km 处。按照前述讨论,Stagno and Frost (2010)通过高温高压实验认为地幔碳酸盐在100 ~150km 时就开始被还原成金刚石(石墨)稳定存在了。在上地幔更深处,Rohrbach et al. (2007)通过高温高压实验研究发现石榴石和辉石在>7GPa 时可以容纳大量的Fe3+,导致在上地幔大于250km 时Fe、Ni 金属可以稳定存在。基于此推论,Rohrbach and Schmidt (2011)进而通过高温高压实验模拟认为俯冲碳酸盐熔体在进入受Fe、Ni 金属控制氧逸度的上地幔时(250km)会变得不稳定,被大量还原成金刚石稳定下来(图4)。以上通过高温高压实验确定的碳酸盐被还原成金刚石/石墨的深度都是建立在全球地幔氧逸度均一的假设基础之上。然而,由于地幔氧逸度空间分布的不均一性,碳酸盐被还原成金刚石的深度会出现一定的变化。譬如携带有大量氧化性物质(H2O 和碳酸盐)的俯冲带进入深部地幔时,就不会在Stagno and Frost (2010)所认为的150km 处发生明显的碳酸盐还原作用。俯冲碳酸盐就可以进入更深的地幔,在到达Rohrbach and Schmidt (2011)认为的250km 时,大量俯冲碳酸盐应该会被具有强还原属性的Fe、Ni 金属还原成金刚石稳定下来。在这种情况下,Dasgupta et al. (2013)认为的碳酸盐化橄榄岩熔体在250 ~200km 处产生的结论应该是合理的。相反,深部地幔稳定的金刚石在地幔柱或者洋中脊岩浆作用过程中,可以被上部地幔中含Fe3+硅酸盐矿物氧化成容易迁移的碳酸岩熔体,逃离地幔,喷出地表(Rohrbach et al.,2011;Stagno et al.,2013)(图4)。

4.6 上地幔C-O-H 流体

在洋中脊玄武岩(MORB)熔融深度(50km),C-O-H 流体主要是以CO2和H2O 为主的氧化性流体。Frost (2006)认为沿着洋中脊绝热线,是不存在任何含水矿物的。根据前述讨论,上地幔150km 深度,地幔碳酸盐开始通过EMOD 反应还原成刚石/石墨稳定存在(Stagno and Frost,2010)。这时,与石墨平衡的上地幔C-O-H 流体相组分及其比例就可以通过流体状态方程进行热力学计算。图5 是基于纯流体计算的上地幔绝热线下C-O-H 流体相组分及其比例。其中氧逸度是假设富集地幔Fe3+/∑Fe 为2%时通过石榴石橄榄岩氧逸度计确定的(Frost and McCammon,2008)。计算表明地幔中含石墨/金刚石的C-O-H 流体中水含量在150km 时达到最大值(图5),这些水可能存在于上地幔名义上不含水矿物(NAMs)中(Beran and Libowitzky,2006;Hirschmann,2006;Skogby,2006)。那么地幔橄榄岩的熔融深度就应该取决于此时地幔中CO2的含量(Dasgupta and Hirschmann,2007)。大于200 km 时,地幔中含石墨/金刚石的C-O-H 流体中CH4和H2含量明显增加(图5)。到上地幔深部(>250km)时,其氧逸度主要受控于与饱和铁、镍金属沉淀曲线(Ballhaus,1995;Rohrbach et al.,2007)。此时地幔中的C-O-H 流体是以CH4和H2O 为主,以及少量的H2(Wood et al.,1990;Ballhaus,1995;Matveev et al.,1997)(图5)。由此可见,真正的洋中脊上地幔的熔融可能取决于地幔深部还原性碳(石墨、金刚石、甲烷)在减压上涌过程中被上部含Fe2O3的矿物氧化成CO2和H2O 的位置。氧化形成的CO2和H2O 可以降低上部地幔熔融温度,形成含碳酸岩的氧化熔体(Taylor and Green,1988;Ballhaus and Frost,1994;Stagno et al.,2013)。通过此过程可以从深部地幔抽取还原性碳C-O-H 流体以氧化性的碳酸盐岩喷出地表。

图5 在1200℃,上地幔条件下通过热力学计算与石墨/金刚石共生C-O-H 流体在图1 所示线1 和线2 氧逸度曲线控制下所出现的物相种类和含量(据Frost and McCammon,2008 修改)Fig. 5 The speciation of the C-O-H fluid phase in equilibrium with graphite/diamond calculated as a function of pressure in the upper mantle along an adiabat with a potential temperature of 1200℃ and at an fO2 defined by curves 1 and 2 in Fig.1 (after Frost and McCammon,2008)

5 俯冲带氧逸度及其对含碳相氧化还原作用

前人岩石学和实验模拟研究俯冲带中主要含碳相(碳酸盐)稳定性,一般都只局限于温度和压力对碳酸岩的稳定性的影响(Zhang and Liou,1994,1996;Martinez et al.,1996;Luth,2001;Sato and Katsura,2001;Shirasaka et al.,2002;Zhang et al.,2003b;Buob et al.,2006;Morlidge et al.,2006;Hammouda et al.,2011)。然而,由于俯冲带碳酸盐中一般都含有C 以及Fe、Mn 等变价元素,使碳酸盐的稳定性也受到了氧逸度这一重要成分变量的影响(Tao et al.,2013)。前述我们讨论了地幔氧逸度对俯冲进入深部地幔含碳相的影响,但是关于俯冲带本身的氧化还原条件及其对含碳相稳定性影响的研究还不是很多(Connolly,1995;Hayes and Waldbauer,2006)。根据前述讨论,前人普遍认为俯冲带中含碳相是以碳酸盐以及H2O 和CO2为主的氧化性C-O-H流体为主(Wood et al.,1990;Parkinson and Arculus,1999;Kelley and Cottrell,2009)。但是Galvez et al. (2013)通过拉曼光谱和碳同位素研究了Alpine Corsica 俯冲带中折返蛇纹岩和富碳酸盐沉积岩接触带中石墨,发现富碳酸盐沉积岩中高度结晶的石墨是通过俯冲带蛇纹岩释放的还原性流体还原富碳酸盐沉积岩而形成的。也就是说在低温俯冲过程中,洋壳超基性岩蛇纹岩化过程可能会导致上部俯冲板片富含大量还原性流体。近来在西南天山俯冲带榴辉岩和泥质片岩中发现了普遍存在的石墨(Lü et al.,2008,2009,2013)和碳氢化合物流体包体(Tao et al.,2015)说明俯冲带含碳相稳定性(碳酸盐)不仅受到温度、压力等因素控制,同时也受到俯冲带氧逸度对含碳相变化及物化属性的影响。俯冲碳酸盐可以被俯冲带本身的低氧逸度被还原成石墨、甲烷等还原性含碳相,从而改变含碳相迁移能力等属性。

鉴于Fe 作为重要的俯冲碳酸盐的成分,而且含铁碳酸盐中Fe 和C 可以通过自身氧化还原作用改变含碳相稳定性的属性。Tao et al. (2013)通过高温高压实验确定了天然菱铁矿(FeCO3)的脱碳反应边界及熔融曲线,与不同地热曲线对比发现纯菱铁矿可以在地幔100km 深度就可以发生脱碳/熔融反应,并通过自身歧化反应产生磁铁矿和石墨/金刚石。同时,Tao et al. (2014)通过岩石学和高温高压实验模拟方法研究了西南天山俯冲含铁白云石的分解结构,发现Fe 可以明显的降低白云石等含铁碳酸盐的高压稳定性。有意思的是,Tao et al. (2015)在西南天山碳酸盐化榴辉岩中发现了典型的含铁碳酸盐在变质流体作用下,通过自身岐化反应产生石墨和磁铁矿的结构,并利用此结构限定出西南天山具有比其它俯冲带低得多的氧逸度(ΔFMQ-1.5)。结合菱铁矿高温高压条件下歧化分解反应产生石墨反应(Tao et al.,2013),提出西南天山变质岩中发现的普遍存在石墨以及少量碳氢化合物很有可能是含铁白云石被变质流体溶解并发生歧化反应形成的。为了验证以上岩石学假设,Tao et al.(2015)通过高温高压实验确定了含铁白云石和水可以通过岐化反应形成石墨、甲烷、乙烷等无机碳氢化合物的事实。通过上述对西南天山俯冲带中不同含碳相(碳酸盐,石墨,碳氢化合物)稳定性及其相关转换关系的岩石学和实验模拟研究,我们认为西南天山俯冲带作为迄今为止发现的两例洋壳超高压变质俯冲带之一(Zhang et al.,2013),其中发育的经历了高压/超高压变质作用的大理岩(Lü et al.,2013)、碳酸盐化泥质变质岩(Zhang et al.,2003)、碳酸盐化榴辉岩(Zhang et al.,2002)都是研究俯冲带深部碳循环的理想天然样品(Tao et al.,2014)。同时西南天山俯冲带中发育的不同含碳相(碳酸盐,石墨,碳氢化合物,含碳C-O-H 流体包裹体)对研究俯冲带氧逸度及其对含碳相转变和迁移属性的影响都就有重要意义(Tao et al.,2015)。

6 结语

通过总结地幔氧逸度的时空分布和演化规律,结合深部碳循环在不同地幔氧逸度条件下的氧化还原作用,我们认为在考虑俯冲带深部碳循环过程时,不能只考虑温度和压力等常规变量,必须要把影响含碳相物相转变和稳定性的氧逸度这一重要成分变量引进来。氧逸度通过影响碳的价态,使含碳相从氧化态的碳酸盐、CO2和CO 到还原态石墨/金刚石、甲烷和铁碳化合物之间转变。同时引起含碳相从易迁移的碳酸盐熔体、CO2、CH4到不易迁移石墨/金刚石和铁碳化合物之间转变。在不同的地球深部环境下,不同的氧逸度条件明显改变了含碳相的存在形式和迁移能力。从俯冲带深部碳循环的角度来说,碳循环在洋壳形成;洋壳水化、碳酸盐化;俯冲脱水、脱碳;深俯冲碳酸盐再循环过程中都会受到不同地质背景下不同氧逸度的影响。地表CO2以氧化性的碳酸盐形式沉淀在俯冲洋壳中,在俯冲进入深部地幔后被还原成石墨/金刚石、碳氢化合物或者铁碳化合物;而深部地幔还原性石墨/金刚石以及CH4流体在洋中脊和地幔柱岩浆作用过程中又可以被氧化成碳酸盐熔体或者CO2流体循环带出地表。从这个角度来说,地球深部碳循环也是伴随碳在不同含碳相之间的氧化还原作用,同时也证明地幔氧逸度引起的含碳相的氧化还原作用是影响俯冲带深部碳循环的非常重要的因素,值得深入谈论和研究。通过对西南天山俯冲带深部碳循环工作的总结,我们认为俯冲带自身氧化还原环境及其对俯冲带深部碳循环的影响是将来研究的重点问题。

致谢 本文是在三位匿名评审人大量建设性意见基础上逐渐成熟定稿的,在此对他们表示由衷的感谢。

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