Interaction between seafloor cold seeps and adjacent hydrothermal activities in the Okinawa Trough
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摘要: 热液和冷泉活动是现代深海环境中两个重要的极端系统,它们均是岩石圈与外部圈层之间进行物质、能量转移和交换的重要途径,它们之间既有显著差异,但也存在很多相似点。一系列调查研究表明,在某些特殊构造单元,热液和冷泉活动可能并不是彼此孤立的,而是在构造地质、生物生态和元素循环上存在某种相互作用或耦合关系。冲绳海槽作为西太平洋一个典型的弧后盆地,发育了繁盛的热液和冷泉活动,是研究这两个海底极端系统相互影响机制的天然实验室。在大量文献调研和野外精细探测结果的基础上,分析了冲绳海槽内相互毗邻的冷泉和热液之间的物质扩散过程及生物地球化学作用,初步建立了两个极端系统内两种不同流体相互作用的概念模型,认识到未来如对两个深海极端环境共生区构造发育特征、地层流体演化、生物群落以及矿物元素组成进行系统分析,将有助于建立更加完善的冷泉-热液两个系统在物质和能量上的耦合关系模型,同时也有助于揭示它们在生物生态之间的沟通融合规律,最终可建立盆地尺度上热液-冷泉区相互作用模式,从而加深对西太平洋甚至全球范围内冷泉-热液两个极端环境系统甚至“流体-固体”耦合的规律性认识。Abstract: As the seafloor extreme environmental systems, both hydrothermal vents and cold seeps are the critical pathways between the lithosphere and exosphere (biosphere, hydrosphere and atmosphere) for transfer and exchange of materials and energies. There are significant differences, but also many similarities between the two systems. Recent investigations show that in some special tectonic units, hydrothermal vents and cold seeps are not isolated from each other, but instead there are some interactions or coupling relationships in terms of tectonic geology, biological ecology and element cyclicity. As a typical back-arc basin in the western Pacific Ocean, there are abundant hydrothermal vents and cold seeps developed in the Okinawa Trough (OT). Therefore, it has become an ideal natural laboratory for studying the two extreme environmental systems and their interactions. On the basis of literature researches and careful field case studies, we investigate the material diffusion process and biogeochemical process between cold seeps and hydrothermal vents adjacent to each other within the trough. A conceptual model is then established for the interactions between the fluids from the two extreme systems. Our results suggest that it would help to establish a more complete model of the coupling relationship between the two systems in the future, if the structural development characteristics, pore fluid evolution, biological communities, and mineral chemistry of the two deep-sea extreme environments are systematically analyzed. Moreover, it will help to reveal the interaction between them in biological ecology and finally establish a model of interaction between hydrothermal vents and cold seeps on the basin scale, so as to better understand the interaction process between cold seeps and hydrothermal systems and even the coupling of " flow-solid” in the Western Pacific Ocean or the whole Earth.
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Keywords:
- hydrothermal activities /
- cold seeps /
- interaction /
- material and energy cycle /
- Okinawa Trough
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自从20世纪70年代首次发现天然气水合物后,天然气水合物作为一种潜在的洁净能源,越来越受到人们的关注。海洋沉积物中的天然气水合物(>600 Gt)是地球上最大的甲烷储库[1-2],通常分布于大陆边缘海底[3]。然而作为一个动态的甲烷储库,在温压条件改变、水合物相平衡状态遭到破坏时,水合物会发生分解,可导致甲烷进入海水在海底形成冷泉。由于甲烷是一种强烈的温室气体,在地质历史时期中由水合物分解大规模释放的甲烷可引起海水酸化和全球气候变化[1,3-4]。
大陆边缘水合物发育区通常存在着富甲烷流体的渗漏现象。一旦甲烷运移到沉积表层遇到由海水扩散下来的硫酸根时,绝大部分甲烷被甲烷厌氧氧化作用(AOM)消耗[5]。这些渗漏的甲烷作为除沉积有机质之外的有机碳源可为微生物群落提供能量,因此,冷泉渗漏区的生物量比非渗漏区的要高,且与渗漏甲烷的强度有一定的联系[6]。同时AOM也会导致孔隙水碱度升高,从而促进沉积物中自生碳酸盐岩的沉淀[7]。冷泉区孔隙水中的DIC和DOC通常具有较高的浓度,并且它们的13C和14C 出现明显亏损,表明这些碳来源于甲烷[8-9]。在甲烷渗漏强烈的地区,大量的甲烷通常是以气泡形式向上快速运移,不能充分被AOM作用所消耗,而进入上覆水体[10-11]。此外,DIC除了在AOM过程中产生外,有机质硫酸根还原作用(OSR)也是DIC的重要来源[12]。这些甲烷和有机质来源的DIC在海水中的释放可能会对局部深海碳库产生影响。
20世纪90年代以来,我国对南海地区进行了多次高分辨率地震调查,通过似海底反射层(BSR)推断南海北部东沙、西沙、台西南和神狐地区广泛发育有高饱和度的天然气水合物[13-24]。对水合物区沉积物和孔隙水的地球化学特征显示存在与甲烷活动有关的亏损13C的碳酸盐岩和硫酸根浓度剖面异常信号[25-30]。最近运用反应运移模型对南海东沙、神狐、西沙等水合物和冷泉发育区海底表层的生物化学反应速率和甲烷通量进行了评估,认识到甲烷活动在上述区域较为活跃,并在部分站位AOM作用主导了沉积表层碳循环过程[23,27,31-38],但对甲烷和有机质参与的微生物主导的各种碳转化通量和最终归宿的定量研究并不多。
神狐海域位于南海北部陆坡,该海域天然气水合物发育区经过水合物钻探航次调查,显示有大面积富甲烷流体活动[39-41]。本文通过对SH-W19-PC和SH-W23-PC两个站位的地球化学分析显示有较强的AOM作用,因此,选取这两个站位重力活塞柱状沉积物孔隙水的地球化学数据进行数值模拟研究,应用反应运移模型定量评估神狐海域两个站位的与甲烷周转有关的生物地球化学反应速率,计算DIC排放到海水中的通量和深部沉积层来源的甲烷通量。通过综合不同碳库之间的转化关系和转化量,评估神狐海域天然气水合物发育区浅表层沉积物碳循环对海洋碳库和海洋环境的潜在影响。
1. 区域概况
南海是西太平洋最大的边缘海,在欧亚板块、印度—澳大利亚板块和太平洋板块共同作用下形成[41]。大中型新生代沉积盆地在南海北部陆坡发育,晚上新世以来北部陆坡陆源有机质大量输入,沉积物快速堆积,为南海北部陆坡气源的形成提供了有利条件[42]。ODP 184航次3个站位均发现了明显富集烃类物质的沉积层,证实南海北部发育一系列含石油和天然气水合物的沉积盆地[43]。
神狐地区位于南海北部陆坡中段,构造上处于珠江口盆地珠Ⅱ坳陷区内[37]。受珠江口新生代沉积盆地形成演化的影响,神狐海域发育了广泛的海底峡谷、中央底辟带和各种类型的底辟构造[44-45],为甲烷流体的运移提供了良好通道。在新近纪神狐海域沉积速率变大,并且有机质含量较高[46]。上述有利的构造背景和沉积环境指示神狐区域具有形成天然气水合物的良好成藏条件[46]。2007年和2015年广州海洋地质调查局在该海区进行了GMGS1和GMGS3两个航次的水合物钻探,均采集到了水合物实物样品[47],同时神狐海域多个站位的沉积物及孔隙水地球化学分析也指示该区域存在富甲烷的流体活动[48]。
2. 采样和方法
2.1 样品采集
SH-W19-PC和SH-W23-PC站位沉积物样品为2016年9月由广州海洋地质调查局“海洋四号”用重力活塞柱状采样器采集。两个站位均位于南海北部陆坡神狐水合物发育区内,水深分别为约1 300 和1 500 m(图1),柱状样全长分别为820 和847 cm。柱样采集后立刻进行孔隙水抽取,使用孔径为0.2 μm的Rhizon采样器采集,采样间隔为10~40 cm,将4 h内获得15~20 mL的孔隙水装在两个10 mL的玻璃瓶内,一瓶内加入10 μL饱和HgCl2用于DIC浓度及其碳同位素分析,另一瓶用高纯浓硝酸酸化处理,用于阴离子和溶解金属元素分析。所有的样品都在4 ℃下保存直至分析。
2.2 孔隙水参数分析测试
硫酸根(SO42−)和钙离子(Ca2+)浓度在中国科学院南海海洋研究所用Dionex ICS-5000+离子色谱仪检测,分析精度优于2%。仪器分析之前SO42− 浓度用超纯水稀释500倍,Ca2+ 浓度用超纯水稀释100倍。阴离子浓度分析时用28 mM的KOH作为淋洗液,IonPac AS11HC作为分析柱。对于阳离子分析,20 mM甲磺酸作为淋洗液,IonPac CS12A作为分析柱。DIC的浓度和同位素的分析在路易斯安那州立大学完成,通过ThermoFinnigan Gas Bench 和ThermoFinnigan Delta V Advantage两个仪器耦合进行检测,利用浓度范围为1 mM至30 mM的标准溶液获得标准曲线。DIC浓度的分析精度优于2‰,同位素精度优于0.1‰,同位素测试结果用V-PDB标准表示。
2.3 数值模型
利用反应运移模型模拟1个固相组分(POC)和4个溶解组分(SO42−、DIC、Ca2+、CH4),用下面两个偏微分方程描述溶解组分(式1)和POC(式2)[49-50]:
$$ \phi \frac{{\partial {{\rm{C}}_a}}}{{\partial t}} = \frac{{\partial \left( {\phi \cdot{D_s}\cdot\frac{{\partial {C_a}}}{{\partial x}}} \right)}}{{\partial x}} - \frac{{\partial \left( {\phi \cdot{v_p}\cdot{C_a}} \right)}}{{\partial x}} + \phi \cdot\mathop \sum R $$ (1) $$ \left( {1 - \phi } \right)\frac{{\partial POC}}{{\partial t}} = - \frac{{\partial \left( {\left( {1 - \phi } \right)\cdot{v_s}\cdot POC} \right)}}{{\partial x}} + \left( {1 - \phi } \right)\cdot\mathop \sum \nolimits R $$ (2) 式中x(cm)为POC或溶解组分所在的沉积深度;t(a)为时间;φ为孔隙度;DS(cm2 a−1)是经曲度矫正过的分子扩散系数;Ca(μmol cm-3)是溶解组分的浓度;POC是沉积物中有机质的百分含量;vp(cm a−1)是孔隙水的埋藏速率;vs(cm a−1)是固体的沉积速率;是在该模型中某一组分变化所考虑的生物化学反应速率总和。在模拟的长度范围内主要考虑的生物化学过程包括:OSR、AOM、产甲烷作用和自生碳酸盐岩沉淀。由于在重力活塞柱状样品采集过程中常导致最顶部的10~20 cm沉积物样品的缺失,故通常不考虑有机质有氧矿化、反硝化作用和铁锰还原作用[51]。
由于孔隙度和沉积速率数据没有实际测得,故参考临近站位的实测值[52]。在海底没有外部流体对流的情况下固相和孔隙水的埋藏速率相同。迂曲度矫正后的分子扩散系数DS由下面方程得到:
$$ {D_s} = \frac{{{D_m}}}{{1 - {\rm{ln}}{{\left( \phi \right)}^2}}} $$ (3) 式中Dm是海水中溶解组分在一定温度盐度和压力下的扩散系数。
有机质矿化为海洋沉积物中的微生物活动提供了能量,在沉积物缺氧层内硫酸根是有机质氧化的主要电子受体[53],OSR过程可被简化为:
$$ 2C{H_2}O{\left( {PO_4^{3 - }} \right)_{{r_p}}} + SO_4^{2 - } \to 2HCO_3^ - + {H_2}S + 2{r_p}PO_4^{3 - } $$ (4) OSR过程中每消耗1 mol硫酸根会氧化2 mol有机碳,同时释放DIC、硫化物和磷酸盐到孔隙水中,rp在是有机磷和有机碳的比例。硫酸根消耗完后有机质产甲烷开始,方程式为:
$$ 2C{H_2}O{\left( {PO_4^{3 - }} \right)_{{r_p}}} \to C{O_2} + C{H_4} + 2{r_p}PO_4^{3 - } $$ (5) 稳态情况下该过程产生等量的甲烷和CO2,自然环境中甲烷主要由有机质发酵和CO2还原作用产生[54]。在模型中OSR和产甲烷过程的反应速率速率取决于POC的总降解速率RPOC(g C g−1a−1)[55]:
$$ {R_{POC}} = \frac{{{K_c}}}{{\left[ {DIC} \right] + \left[ {C{H_4}} \right] + {K_c}}}\cdot\left( {0.16\cdot{{\left( {{a_0} + \frac{x}{w}} \right)}^{ - 0.95}}} \right)\cdot POC $$ (6) 式中KC是孔隙水中DIC和CH4积累对POC降解的抑制系数;[DIC]和[CH4]分别代表DIC和CH4的浓度;a0是有机质初始年龄,反映其活性程度[56]。OSR速率(RSR,μmol cm−3a−1,以SO42−计),产甲烷速率(RMG,μmol cm−3a−1,以CH4计)和POC产DIC速率(RDP,μmol cm−3a−1,以C计)表达式如下:
$$ {R_{OSR}} = 0.5\cdot\frac{{{\rho _s}\cdot\left( {1 - \phi } \right)\cdot{{10}^6}}}{{M{W_c}\cdot\phi }}\cdot\frac{{\left[ {SO_4^{2 - }} \right]}}{{\left[ {SO_4^{2 - }} \right] + {K_{SO_4^{2 - }}}}}\cdot{R_{POC}} $$ (7) $$ {R_{MG}} = 0.5\cdot\frac{{{\rho _s}\cdot\left( {1 - \phi } \right)\cdot{{10}^6}}}{{M{W_c}\cdot\phi }}\cdot\frac{{{K_{SO_4^{2 - }}}}}{{\left[ {SO_4^{2 - }} \right] + {K_{SO_4^{2 - }}}}}\cdot{R_{POC}} $$ (8) $$ {R_{DP}} = \frac{{{\rho _s}\cdot\left( {1 - \phi } \right)\cdot{{10}^6}}}{{M{W_c}\cdot\phi }}\cdot{R_{POC}} - {R_{MG}} $$ (9) 式中ρs是沉积物干重的密度,MWc是C的相对原子质量,KSO42−是在低SO42−浓度时的SO42−还原作用的抑制系数。
AOM反应总的方程式为:
$$ C{H_4} + SO_4^{2 - } \to HCO_3^ - + H{S^ - } + {H_2}O $$ (10) AOM反应速率(RAOM)可由双分子动力学方程表示[55]:
$$ {R_{AOM}} = {k_{AOM}}\cdot\left[ {SO_4^{2 - }} \right]\cdot\left[ {C{H_4}} \right] $$ (11) 式中kAOM是AOM反应速率常数,Ca2+碳酸盐岩中的沉淀(或形成)速率(RPPT),可简化为由模型模拟出的浓度和实测浓度之差计算:
$$ {R_{PPT}} = {K_{PPT}}\cdot\left( {\left[ {C{a^{2 + }}} \right] - {{\left[ {C{a^{2 + }}} \right]}_T}} \right) $$ (12) 式中KPPT是一级速率常数,[Ca2+]和[Ca2+]T分别是模拟和实测站位(SH-W19-PC和SH-W23-PC)各层位孔隙水Ca2+浓度。
为了求解这个模型,设定模型上边界如表1所示,SO42−、Ca2+、DIC和CH4的浓度为固定值(狄利克雷型边界条件),上边界值为海水中各组分的浓度,由于没有实测甲烷数据,故甲烷浓度设为0;POC的上边界设为通量值(罗宾型边界条件);模型的下边界除了甲烷浓度外,各组分的浓度梯度都设为0(纽曼型边界条件);下边界甲烷浓度是通过较好的拟合SO42−浓度剖面获得。为了保证有机质有较高的降解程度,模型模拟长度为20 m,在不均一的网格内(由表层到深层分辨率逐渐降低)利用有限差分和直线法将式(1)、式(2)两个偏微分方程离散化,最后用MATHMATICA V.8.0中NDSolve命令结合反应速率方程求解并计算出各个反应过程的反应速率和各组分通量及转化量。
表 1 SH-W19-PC和SH-W23-PC站位模型参数及边界条件Table 1. Model parameters and boundary conditions at sites SH-W19-PC and SH-W23-PC参数 SH-W19-PC SH-W23-PC 单位 温度(T) 5 5 ℃ 盐度(S) 33.5 33.5 % 压力(P) 105 105.1 bar 干燥固体密度(ρS) 2.5 2.5 g/cm3 沉积速率(ω)a 0.033 0.033 cm/a 沉积物-水界面孔隙度(φ0)b 0.7 0.7 — POC初始年龄(a0)c 40 40 ka SO42−在海水中的扩散系数c 191 191 cm2/a CH4在海水中的扩散系数c 294 294 cm2/a DIC在海水中的扩散系数d 203 203 cm2/a Ca2+在海水中的扩散系数c 142 142 cm2/a POC抑制系数(KC)e 35 35 mmol/L POC降解米氏常数(KSO42-)e 0.1 0.1 mmol/L AOM的动力系数(KAOM)e 1 1 cm3/(a·mmol) SO42−的上边界条件 28.2 28.2 mmol/L DIC的上边界条件 2.2 2.3 mmol/L Ca2+的上边界条件 11 11 mmol/L a为南海北部6个ODP184钻孔的平均值[52],b据Wang 等[52],c据Hu Y等[33],d基于碳酸氢根离子的分子扩散系数[27],e据Wallman 等[50]。 3. 结果
3.1 孔隙水地球化学
图2和图3为南海北部神狐区域SH-W19-PC和SH-W23-PC两个站位的SO42−、Ca2+、DIC和δ13CDIC随深度变化的剖面图。SH-W19-PC相较SH-W23-PC站位的SO42−浓度降低更快。SH-W19-PC站位由表层到760 cm处,SO42−浓度从28.2 mM减小至0。Ca2+ 浓度从表层的12.1 mM减小到底部3.7 mM。DIC浓度随深度明显增加,变化范围为2.2~35 mM。同时δ13CDIC在深度剖面上逐渐减小,在DIC浓度最大的深度720 cm处达到最小值(−25.1‰)。SH-W23-PC站位SO42−和Ca2+浓度在深度上几乎都显示线性减小趋势,SO42−浓度变化由表层28.2 mM减小至800 cm处的7.7 mM。Ca浓度从表层10.7 mM减小至底部4.7 mM。DIC浓度随深度逐渐增大而δ13CDIC随深度减小,在800 cm处δ13CDIC达到最小值−23.9‰。
图 2 SH-W19-PC和SH-W23-PC站位SO42−、Ca2+、DIC、CH4浓度在深度剖面上实测值和模型拟合结果 (红色点代表实测浓度值,黑色线为模型拟合曲线)Figure 2. Measured values and simulate depth profiles of core SH-W19-PC and SH-W23-PC.Down-depth concentration of SO42−, DIC, Ca2+and CH4 are shown (red dots represent measured concentration values, black lines are model fitting curves)3.2 模型计算结果
两个站位模拟孔隙水Ca2+、DIC、SO42−剖面都较好的拟合了实际数据。两个站位POC的初始年龄a0和有机质含量参考附近站位值[33],SH-W19-PC和SH-W23-PC两个站位模拟有机质降解速率分别为14.4和15.4 mmol∙m−2a−1(表2)。SH-W19-PC的OSR反应速率(9.2 mmol∙m−2a−1)小于SH-W23-PC(11.6 mmol∙m−2a−1),两个站位的产甲烷速率相当。SH-W19-PC站位下部的甲烷通量(25.9 mmol∙m−2a−1)明显高于SH-W23-PC(18.4 mmol∙m−2a−1),同时SH-W19-PC的AOM反应速率(28.5 mmol∙m−2a−1)也比SH-W23-PC(20.2 mmol∙m−2a−1)要高,这主要是由于甲烷下边界不同而导致的通量不同。利用Ca2+浓度变化计算出两个站位的自生碳酸盐岩沉淀速率分别为9.3和6.6 mmol∙m−2a−1。
4. 讨论
4.1 甲烷周转与DIC源汇
在海洋缺氧沉积物中,甲烷的厌氧氧化过程主要发生在硫酸盐甲烷转换带(SMTZ)内,SMTZ内的甲烷可能为多种来源(生物成因、热成因、水合物分解等),硫酸盐主要是海水中的硫酸根在浓度梯度作用下扩散至该带内[57-58]。SMTZ的深度受多种因素的控制,其中甲烷通量的大小对该带的影响显著[59-60]。随着甲烷通量的增加,SMTZ深度逐渐变浅,硫酸盐浓度剖面呈线性变化并且斜率随之增加[61]。对于全球陆坡区,SMTZ深度的平均深度为1 280 cm[61]。在水深为1 500 m附近的神狐海域两个站位,它们的SMTZ深度明显小于平均值,可能表明具有较大的甲烷通量。两个站位孔隙水DIC浓度随深度急剧增加,同时δ13CDIC逐渐亏损至−25‰附近,说明孔隙水DIC库中有AOM来源13C亏损的DIC加入[62]。比较两个站位发现SH-W19-PC的硫酸根浓度梯度大于SH-W23-PC,SMTZ深度小于SH-W23-PC,表明SH-W19-PC站位下部甲烷通量可能更大[60]。数值模拟结果证实了上述推断,SW-19站位下部来源的甲烷通量(25.9 mmol∙m−2a−1)大于SW-23(18.4 mmol∙m−2a−1)。同时两个站位的有机质产甲烷速率分别为2.6和1.9 mmol∙m−2a−1。而且模型计算的AOM反应速率SW-19站位也高于SW-23站位(分别为28.5和20.2 mmol∙m−2a−1)(表2),总体上与先前该海区的研究结果相近(20.9 mmol∙m−2a−1)[63],但均低于东沙水合物发育区的值[32-33]。
自生碳酸盐岩的沉淀主要受孔隙水总碱度和Ca2+浓度的控制。由于孔隙水碱度随AOM作用强度的增加而增加,所以自生碳酸盐岩沉淀速率SH-W19-PC站位(9.3 mmol∙m−2a−1)高于SH-W23-PC站位(6.6 mmol∙m−2a−1),但略小于神狐海区的另一冷泉活动站位(10.43 mmol∙m−2a−1)[62]。
通常来讲,正常沉积物孔隙水DIC的源主要有OSR、AOM和ME过程,DIC的汇主要包括自生碳酸盐岩沉淀、向海底扩散到进入海水中和沉积埋藏部分[63]。以上6个过程直接影响SH-W19-PC和SH-W23-PC两个站位的DIC库。对于DIC的来源,两个站位AOM作用产生的DIC占总DIC产量的比例分别为70.7%和60%,每个过程对DIC产量的贡献按AOM,OSR和ME的顺序依次减少。这与南海北部东沙和西沙海域相一致[59],但是该顺序又和郁陵盆地有所不同(AOM>ME>OSR),这可能是因为有机质活性和模拟深度不同所导致[64]。对于DIC的汇,SH-W19-PC和SH-W23-PC两个站位的DIC向海底排放的通量分别为24.7和20.7 mmol∙m−2a−1,均占DIC汇的约60%。另一部分DIC在沉积物中或在海底表层与Ca2+结合形成自生碳酸盐岩沉淀。研究区两个站位形成碳酸盐的DIC分别占总DIC汇的23%和20%,较小部分的剩余的DIC随沉积埋藏保存在孔隙水中。
4.2 浅表层沉积物碳循环对海洋碳库的潜在影响
作为海洋碳循环的重要组成部分,沉积表层有机碳埋藏、有机质降解和甲烷氧化作用共同调节了深海碳库、沉积层碳库和孔隙水碳库之间的平衡,并持续影响海洋环境[65]。由河流带来和海表真光层形成的颗粒有机碳(POC)在沉降过程中被微生物降解,到达水深大于1 000 m海底的POC只占表层初级生产力的1%~3%[66],从神狐SH-W19-PC和SH-W23-PC两个站位的数值模拟结果看(图4),海底沉积表层POC通量为100 mmol·m−2a−1,在模型下边界POC的埋藏通量分别为85.6和84.6 mmol·m−2a−1,计算获得这两个站位有机质降解速率为14.4和15.4 mmol·m−2a−1(表2)。在模型中这部分有机碳被OSR和产甲烷作用利用,这些原位产生的甲烷与下部来源的甲烷在SMTZ内消耗,产生大量DIC,进而扩散到上覆海水中。由于SH-W19-PC和SH-W23-PC站位甲烷活动强度较弱,甲烷都是以扩散形式向上运移,不存在甲烷气泡的运移方式,所以全部甲烷在近海底沉积物孔隙水中被消耗,进入海水中的甲烷通量接近0。通过对比两个站位碳转化关系可知,DIC是沉积表层碳循环的重要载体。随着甲烷通量的增加,由沉积物向海水中释放的DIC也随之增加,尽管计算结果比海岸带和大陆架的DIC通量小很多[67],但这些DIC可能对冷泉区底部水的化学性质产生一定影响。
根据DIC碳同位素的不同,通常可以对不同来源DIC进行划分。海水中的δ13CDIC为0,海洋真光层POC来源δ13CDIC约为−20‰,陆源有机质来源δ13CDIC不超过−25‰,而甲烷来源δ13CDIC均小于−40‰[68]。研究区两个站位δ13CDIC在柱样下部都达到−25‰附近,南海沉积有机质的δ13CDIC值约为−20‰[62,69],在模型中有机质降解过程包括OSR和ME两个过程,OSR过程不考虑碳同位素的分馏效应。ME的产物甲烷和DIC发生碳同位素分馏,但是这部分甲烷通过AOM作用产生的DIC又与之前的DIC混合,所以认为有机质降解来源的DIC碳同位素值与有机质一致[33]。那么神狐海区沉积物有机质来源的DIC和海水中的DIC混合不可能达到−25‰。通过数值模拟计算得出两个站位存在不同程度的AOM作用,可以认为两个站位孔隙水DIC是3个端元混合的结果。模拟结果显示AOM作用产生的DIC占总DIC产量的比例最大。从沉积物向海水中释放的DIC通量来看,即使所有的有机质来源的DIC全部被排放到海水中,在这两个站位也仅占沉积物—水界面DIC通量的47%和65%,所以甲烷来源的DIC占沉积物—水界面DIC通量的较大比例。本研究区两个站位甲烷来源的DIC及其亏损13C的特征所造成的潜在影响可能主要包括:①在沉积物中被化能自养生物利用,导致生物体碳同位素异常[70];②AOM作用产生的DIC导致孔隙水碱度升高,促进亏损13C的自生碳酸盐岩的形成[71-72];③向海水中排放亏损13C的DIC,可能对冷泉区域局部海底碳库的组分造成一定的影响。
此外,在地质历史时期由海底沉积物向海洋排放的甲烷对于调节海洋碳库和全球气候变化起到了重要作用[61]。由本文的模拟研究结果可知,向海底排放的甲烷有一部分通过AOM作用以DIC形式进入了海洋,因此,推测当地质历史时期海底发生大规模甲烷渗漏活动时期,甲烷的最终归宿可能是以DIC形式大量进入海水中,从而深刻影响海洋无机碳库。
DOC作为碳循环过程中的重要组成部分,目前对于冷泉区沉积物向海水中释放的DOC已有少量研究[73-75],研究表明随着海底甲烷通量的增加,以甲烷作为能源的微生物在代谢过程中将产生更多的DOC,在一些地区甲烷来源的DOC可达到沉积物向海水释放DOC通量的28%[9],这些DOC的活性成分可以被微生物利用进而转化为DIC或惰性DOC。由于缺乏实测DOC数据,无法对SH-W19-PC和SH-W23-PC两个站位孔隙水DOC的活性及通量进行判断,但是甲烷和有机质来源的DOC与DIC的转化是不能忽视的。因此,通过冷泉区甲烷来源的DOC的研究将有助于增强对深海DOC参与的生物地球化学循环的认识[9]。
5. 结论
本文利用反应运移模型对南海北部神狐海域SH-W19-PC和SH-W23-PC两个站位沉积物浅表层物理化学过程进行模拟,定量计算出碳循环各个过程的反应速率和转化量。计算结果显示两个站位平均AOM反应速率(24.3 mmol·m−2a−1)明显高于有机质降解速率(14.9 mmol·m−2a−1),OSR和ME过程产生的DIC量约为AOM作用产生DIC的一半。两个站位DIC汇的主要形式是由沉积物释放到海水中,沉积物—水界面DIC通量分别为(24.7和20.7 mmol·m−2a−1),少部分以自生碳酸盐岩的形式保存于沉积物中。因此,在研究区两个站位甲烷活动主导了沉积物浅表层的碳循环过程,大量甲烷来源的DIC释放到海水中。这些由沉积物向海水中排放的甲烷来源亏损13C的DIC,在地质历史时期海底甲烷活动强烈时期可能会深刻影响深海碳库组成和生态环境。
致谢:非常感谢海洋四号机组人员在海上采样的帮助。
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图 1 世界海洋中已发现热液喷口(蓝色圆点)和可能冷泉发育区(粉色区域,以水合物稳定带代表)的叠置图
① 东海冲绳海槽,② 马里亚纳海沟和日本海沟,③ 西南太平洋(以劳盆地和马努斯海盆为典型),④ 东北印度洋脊,⑤ 东地中海,⑥ 北极加克超慢速扩张洋脊,⑦ 北大西洋南缘,⑧ 东北太平洋边缘(底图据Sun et al[12]修改)
Figure 1. Geological settings where hydrothermal fields and cold seep systems coexist in the global oceans
① Okinawa Trough, East China Sea, ②Mariana and Japan Trench, ③ Extending basins along the Southwest Pacific Ocean (e.g., Manus basin and Lau Basin), ④ Northeast Indian Ridge (Carlsberg Ridge), ⑤ Eastern Mediterranean, ⑥ Arctic Gakkel ultraslow spreading ocean ridge, ⑦ South margin of North Atlantic, ⑧ Northeast Pacific margin. Base map is modified after Sun et al[12]
图 2 典型泥火山型冷泉(A)和洋中脊热液喷口(B)系统剖面图
A图中,AOM:甲烷厌氧氧化反应,GHSZ:天然气水合物稳定带,MDAC:甲烷厌氧氧化来源的自生碳酸盐岩。A图修改自Ceramicola et al.[25];B图修改自German et al.[26]
Figure 2. Typical profiles of seafloor mud volcanic cold seeps (A) and hydrothermal vents (B) on a mid-ocean ridge
In subfigure A, AOM:anaerobic oxidation of methane, GHSZ:gas hydrate stability zone,MDAC:methane-derived authigenic carbonate. Subfigure A is modified from Ceramicola et al.[25] and subfigure B from German et al.[26]
图 3 加利福尼亚湾内瓜伊马斯盆地扩张中心附近火成岩与冷泉之间的关系[51]
A. 多道地震剖面指示了岩床与气体活动(冷泉)的关系;B. 岩床与上部反射紊乱区域(绿色),显示沉积物横向尖灭岩床侵位后的上超现象;C. 年轻岩床上部存在浅层气现象,海底探测到冷泉生物群落;D. 裂谷轴线处的丘体下方存在岩床;E. 浊积层下部存在碟形岩床
Figure 3. Relationship between seafloor cold seeps and nearby igneous sills across the northern Guaymas spreading segment, according to seismic observations[51]
A. Time-migrated MCS section with the amplitude coloured for large values, which tend to indicate sills, gas or turbiditic strata. The green rectangles indicate the bottom panels of depth-migrated detail; B. Sills with overlying disturbed region (green), lateral termination of disturbed region, thickness of post-intrusion sediments, and onlap onto the post-intrusion sea floor indicated; C. Shallow gas above interpreted young sill; a seafloor community is located above this feature; D. Sills beneath mound within axial graben; E. Saucer-shaped sills beneath turbiditic sediments
图 4 冲绳海槽地理位置上相邻的热液喷口和冷泉区的极端生态群落对比
A. 南奄西热液区的管状蠕虫,来自Watanabe, et al. (2015)[56];B. 南奄西热液区的贻贝和毛瓷蟹,来自Watanabe, et al. (2015)[56];C. GT-D1冷泉区的管状蠕虫和巨蛤,ROV摄像,2017年张謇号调查航次;D. GT-D1冷泉区巨蛤床,ROV摄像,2017年张謇号调查航次
Figure 4. Comparison of extreme ecological communities between adjacent hydrothermal vents and cold seeps in the Okinawa Trough
A. Tubeworm clump in the Minami-Ensei Knoll, c.f. Watanabe et al. (2015)[56]; B. Typical rocky fauna in the Minami-Ensei Knoll, c.f. Watanabe et al. (2015)[56]; C. Tubeworm and clam from the GT-D1 cold seep site. Image by ROV Beaver during the integrated environmental and geological expedition of R/V Zhangjian in 2017; D. Clam bed observed in the GT-D1 cold seep site. Image by ROV FCV3000 during the integrated environmental and geological expedition of R/V Zhangjian in 2018
图 6 冲绳海槽热液—冷泉两个系统流体相互作用概念模型
具体内容见正文。OZ:氧化带,SMTZ:硫酸盐—甲烷过渡带,SDZ:硫酸盐亏损带,S-AOM:硫酸盐还原驱动的甲烷厌氧氧化作用,Fe-AOM:铁氧化物还原驱动的甲烷厌氧氧化作用
Figure 6. Conceptual model of fluid interaction between hydrothermal vents and adjacent cold seeps on the western slope of the OT, see the context for details
OZ:oxidative zone; SMTZ:sulfate-methane transition zone; SDZ:sulfate-depleted zone; S-AOM:sulfate reduction-driven AOM; Fe-AOM:Fe oxide reduction driven AOM. Not to scale
表 1 海底冷泉和热液系统之间的异、同特征比较
Table 1 Comparison of the characteristics of seafloor cold seeps and hydrothermal systems
相似点 都具有重要的资源效应 均支持化能自养生物群落 具有相似的环境效应(圈层之间物质和能量交换) 不同点 流体来源和成因机制不同 地质构造和发育位置不同 生化反应和元素循环不同 -
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