Geochemical characteristics of sediment pore water in Haima area of the South China Sea: An indication of cold seeps
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摘要:
海马冷泉位于南海琼东南海域，是南海迄今发现的两个活动冷泉之一。我们对海马冷泉Rov2和PC3站位两个活塞重力柱沉积物孔隙水的阴阳离子、溶解无机碳（DIC）及其碳同位素组成和Sr、Ba含量等进行了分析。结果表明，两个站位孔隙水DIC含量（Rov2和PC3最大DIC含量分别为27.4 、8.5 mM）和δ13CDIC（Rov2和PC3站位最低值分别为−54.63‰和−48.93‰）具有明显的镜像关系。结合孔隙水硫酸盐浓度的变化特征，Rov2和PC3站位的硫酸盐-甲烷界面（SMI）分别位于约485和410 cm。通过模拟估算，Rov2和PC3站位向上甲烷通量分别为67.4和97.2 mol·m−2·ka−1，较浅的SMI深度与相对较高的甲烷通量相一致。SMI附近极低的孔隙水δ13CDIC值指示了AOM作用的发生及其对DIC的贡献。在Rov2站位，自生碳酸盐矿物以高镁方解石为主，阳离子Ca2+、Mg2+和Sr2+含量随深度增加并表现出与SO42−阴离子含量相似的变化特征。在SMI附近，随着SO42−的消耗、有机质的矿化将大量的Ba2+和PO43−释放进入孔隙水。因此，冷泉孔隙水地球化学特征的变化能帮助我们有效识别渗漏活动过程，对AOM作用下物质的迁移与转化具有重要的指示意义。
Abstract:The Haima cold seeps are located in the southeastern part of Qiongdongnan Basin, which is one of the two active cold seeps found in the South China Sea. We analyzed the contents of anions and cations, dissolved inorganic carbon (DIC) and its carbon isotopic composition, and Sr and Ba contents of sediment pore water in two piston gravity columns at the Rov2 and PC3 cores in Haima Cold Seeps. Results show that the DIC contents and δ13CDIC values of pore water in the two cores had a significant "mirror" relationship. With the increase of depth, the DIC contents of the two cores gradually increased (Maximum DIC content of Rov2 and PC3: 27.4 and 8.5 mM, respectively). In contrast, the δ13CDIC values had a negative excursion (Minimum values for the two cores: −54.63‰ and −48.93‰, respectively). Combined with the sulfate depth profile characteristics of pore water, the sulfate-methane interface (SMI) in Rov2 and PC3 cores was located at ~485 and ~410 cm, respectively. The upward methane fluxes in Rov2 and PC3 cores were estimated to be 67.4 and 97.2 mol m−2 ka−1, respectively. The very low δ13CDIC values in pore water near SMI are indicative of the occurrence of AOM (anaerobic oxidation of methane) interaction and its contribution to DIC. In Rov2 core, authigenic carbonate minerals are dominated by high-Mg calcite, and the Ca2+, Mg2+ and Sr2+ showed similar trends to those of SO42−. Near the SMI, with the depletion of SO42−, the mineralization of organic matter released large amounts of Ba2+ and PO43− into the pore water. The geochemical characteristics of pore water could help us effectively identify the early diagenesis in seepage activity area, and are indicative of migration and transformation of materials under the influence of AOM.
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Key words:
- anaerobic oxidation of methane /
- cold seeps /
- pore water /
- South China Sea
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冷泉是指广泛发育于大陆边缘、来自海底沉积界面之下、与海水温度相近、以甲烷等碳氢化合物为主的流体在海底的渗漏活动[1-3]。流体渗漏过程可能诱导海底地质灾害并引发全球气候变化[4-5]。因此，了解海底冷泉的活动规律具有十分重要的意义。
作为重要的温室气体，甲烷的温室效应是同等重量CO2作用的20倍之多[6]。但研究表明，从海洋向大气中释放的甲烷含量仅为全球大气甲烷总量的1%～5%[7-8]，这主要是由于渗漏流体在向上迁移的过程中发生了强烈的甲烷厌氧氧化作用（anaerobic oxidation methane, AOM），将绝大部分甲烷消耗，从而有效地过滤了进入水体的甲烷[8-9]。
CH4+SO2−4→HCO−3+HS−+H2O(甲烷厌氧氧化，AOM) (1) 2CH2O+SO2−4→2HCO−3+H2S(有机碎屑硫还原，OSR) (2) 根据反应式（1），硫酸盐（SO42−）驱动的甲烷厌氧氧化产生大量的溶解无机碳（DIC）与硫化氢（H2S），使孔隙水化学环境发生变化并显著影响早期成岩作用。由于海洋中有机碎屑硫还原（OSR）也通过消耗硫酸盐产生DIC与H2S（反应式（2）[10]），因此充分认识AOM的相对贡献是十分必要的[11-16]。
在冷泉系统中，AOM作用主要发生于硫酸盐-甲烷转换界面（SMTZ），SMTZ的深浅则受冷泉流体渗漏强度的影响[17-19]。由于冷泉系统中的甲烷通常为生物成因或热成因，AOM作用生成产物DIC促进了沉积物中无机碳库的改变 [2, 20-21]。相对于海水和OSR产生的DIC，甲烷厌氧氧化作用所产生的DIC具有较轻的碳同位素特征，亏损13C的孔隙水DIC代表了AOM过程轻碳物质加入了孔隙水DIC库[12, 22]。因此，孔隙水中DIC及碳同位素组成也被作为识别碳源和理解相关生物地球化学过程的重要参数[12, 23-25]。随着大量DIC生成和碱度增加，使得自生碳酸盐矿物发生沉淀[26-27]。受晶格组成的差异控制，不同冷泉碳酸盐矿物所吸收的离子有所不同：具有大的晶格空间的文石倾向于吸收离子半径大于Ca2+的元素，而方解石则更倾向于吸收离子半径小于Ca2+的元素，因此，Sr2+/Ca2+与Mg2+/Ca2+比也常被用来辨别冷泉体系中碳酸盐矿物类型[28-29]。作为重要的汇，碳酸盐、黄铁矿等自生矿物沉淀将活动冷泉区中的碳、硫等元素固定，影响着元素的循环过程[30-31]。AOM作用同时也使孔隙水中Br−、I−、PO43−等生物相关元素表现出一定的异常[32-33]。研究也表明，冷泉活动区AOM作用重塑了孔隙水中微量元素的分布，如：溶解Ba2+会随甲烷流体一起释放，显著影响区域性海水和表层沉积物中Ba的收支平衡[18, 34-37]。因此，SMI附近所形成的“Ba锋”成为判别冷泉活动的指标之一[35, 38]。
我国对于海底活动冷泉的研究开展相对较晚，尤其是南海北部陆坡西北部的琼东南海域仍处于刚刚起步的阶段。直到2015年3月，广州海洋地质调查局在研究区发现与天然气水合物赋存相关的活动冷泉—海马冷泉[39]。目前海洋地质学领域对于海马活动冷泉的地球化学研究仍主要集中在沉积物与孔隙水地球化学、海底自生冷泉碳酸盐及黄铁矿等方面[39-42]，对早期成岩作用中AOM过程的识别、AOM作用下孔隙水中各物质组分的迁移与转化的认识还不够。本研究中，我们对琼东南海域海马冷泉区Rov2和PC3两个沉积柱孔隙水进行了详细的地球化学分析，有助于我们更加清晰地认识冷泉区独特的生物地球化学过程、全面了解冷泉区沉积物与孔隙水之间的相互作用及元素的迁移转化。
1. 地质背景
本次研究区位于南海北部陆坡西北部的海马冷泉区边缘带（图1）。海马冷泉区域水深1370～1390 m[45-46]，总面积约618 km2，其中已发现的活动冷泉区面积约350 km2[39]。
图 1. 研究区域位置Figure 1. The location the study areaa: Bathymetry map of the northern South China Sea region [43]. The dashed red line marks the boundary of the Qiongdongnan Basin. Bottom simulating reflector (BSR) distribution is mapped in the study area (shaded area) [44]. Solid orange line is the “Haima”cold seepage area. The study area (red rectangle) is situated at the southern uplift of the Qiongdongnan Basin; b: bathymetric map of two sampling sitesin the study area.海马冷泉区位于南海琼东南盆地南部，盆地构造上属新生代大陆边缘裂谷形成的复合盆地[47]，发育富有机质的第四纪沉积和新近系上新统海相泥岩[48]，为两套含油气系统沉积序列，具有良好的生烃环境。第四纪以来，较高的沉积速率（最高可达1.2 mm/a[44, 49]）、适宜的温压条件[50-51]，为水合物的形成创造了有利条件。同时，盆地内构造发育，与甲烷渗漏相关的羽流、空白带、气烟囱、麻坑等被相继报道[43, 52-56]，其为甲烷运移提供了良好的通道。丰富的气源及有利的生烃条件，外加良好的构造地质特征，使琼东南盆地成为巨大的水合物资源远景区。广州海洋地质调查局在盆地冷泉区先后开展了现场水合物勘探，利用水下可视采样技术，实地观测到了活跃的冷泉羽流、大型无脊椎动物群以及广泛分布的烟囱、结壳状和块状冷泉碳酸盐岩等[46, 57-59]，并多次在钻孔和表层沉积物中采集到块状天然气水合物样品[39]。据粗略估计，盆地内水合物远景总量可达9.91×109 m3，天然气远景预期高达1.65×1012 m3[60]。
2. 样品采集与分析
2.1 样品现场采集
样品来源于2022年5月广州海洋地质调查局“海洋六号”水合物调查航次获得的琼东南海域Rov2和PC3站位两根大型活塞重力柱样，岩芯长分别为7.15、5.1 m，两站位水深约1400 m。
沉积物柱状样取到甲板后立即进行现场孔隙水提取。按照从顶到底10～20 cm间隔在PVC管上打孔，将Rhizon采样器一端插入沉积物中，另一端用带有0.45 μm滤膜的20 mL注射器抽真空，持续2～4 h采集10～15 mL孔隙水，共获得水样65个。值得注意的是，在Rov2柱样200 cm深度、PC3柱样335 cm深度以下孔隙水样品呈淡黄色到黄色，并具有强烈的臭鸡蛋气味，为H2S气体。所有采集的水样在4 ℃下保存。
2.2 地球化学分析方法
样品运回至实验室后，随即开展地球化学分析，测试项目包括阴阳离子、微量元素含量、溶解无机碳（DIC）的碳同位素组成（δ13C）等。
SO42−、PO43−、Br−含量测试采用美国戴安ICS-2100型离子色谱仪，分析柱为Dionex IonPac AS18（4 mm×250 mm），淋洗液为纯水，流速1.0 mL/min；Na+、Ca2+、Mg2+含量采用瑞士万通1-930-2360型离子色谱仪进行测试，配备Metrosep C4-150/4.0阳离子分析柱，淋洗液为1.6 mmol/L硝酸 + 0.75 mmol/L吡啶二羧酸，流速0.9 mL/min。所有孔隙水样品用超纯水按照1∶100比例稀释。对标准海水的重复测量表明阴阳离子标准偏差均＜2%。
孔隙水Sr2+、Ba2+含量应用等离子体发射光谱仪进行测定，仪器为美国Thermo Fisher公司生产的ICAP-6300。测定前用2%的HNO3介质将孔隙水稀释至适当倍数，测试精度高于0.1 μg/L。
孔隙水DIC及δ13C值分析参照前人连续流质谱法[61-62]进行，所用仪器为美国Thermo Fisher公司生产的MAT 253型GasBench II-IRMS仪。测试时，先向圆底顶空瓶中加入10滴无水磷酸并经He气排空处理（流速100 mL/min，时间9 min），后用1 mL注射器向其内注入约0.3 mL孔隙水样品，涡旋均匀后室温放置一夜使之充分反应。利用70 ℃炉温经毛细管色谱柱（Pora Plot Q，30 m × 0.32 mm × 20 μm，美国Agilent公司）将生成的CO2分离在质谱仪上测定δ13C。δ13CDIC值以国际标准VPDB（Vienna Peedee Belemnite）作为参考，分析精度＜0.10‰。同时，将产生的CO2信号强度用NaHCO3标准溶液的DIC浓度曲线进行校正，计算孔隙水DIC含量，重复分析的精度小于0.5%。
2.3 扩散通量计算
假定为稳态条件[63]，利用Fick第一定律（等式3和4）计算SO42−、Ca2+、Mg2+、Ba2+的扩散通量，公式为：
J=−ΦDsdC/dx (3) Ds=D0/[1−ln(Φ2)] (4) 其中，J为扩散通量（mol·m−2·ka−1或mmol·m−2·ka−1）；Φ为沉积物孔隙度；Ds为各离子沉积物中扩散系数（m2·s−1），D0为海水中各离子自由扩散系数（m2·s−1），C为孔隙水中各离子浓度，x为沉积物深度（m）。研究站位底水温度为3℃[64]，孔隙度为0.69[65]。依据0℃和5℃时各溶质的扩散系数和温度依赖方程，得到3℃时SO42−、Ca2+、Mg2+和溶解Ba2+的海水扩散系数分别为5.288×10−10、3.924×10−10、3.662×10−10、4.352×10−10 m2·s−1[63, 65]。
3. 测试结果
3.1 孔隙水DIC和δ13CDIC值
Rov2和PC3两站位孔隙水DIC和δ13CDIC具有明显的镜像关系（图2）：两站位上部，DIC含量随深度增加而增加，δ13CDIC值随着深度的增加逐渐降低，尤其是在约245、235 cm之后变化趋势变陡。在Rov2岩芯440 cm、PC3岩芯410 cm处，孔隙水DIC和δ13CDIC值达到极值，Rov2岩芯DIC含量最大为27.442 mM，δ13CDIC值最小为−54.63‰；PC3岩芯DIC含量最大为8.507 mM，δ13CDIC值最小为−48.93‰。
3.2 孔隙水中阳离子含量变化
Rov2和PC3站位沉积物孔隙水Na+、K+离子含量整体上保持稳定（图3）。Ca2+、Mg2+、Sr2+阳离子含量具有相似的变化特征，随着深度增加，表现出逐渐降低的趋势。其中，Ca2+的变化最为明显，在上部Ca2+浓度降低的程度相对比较缓慢，约245、235 cm之下则迅速降低，均在岩芯底部浓度降至最低值，与SO42−含量变化趋势相一致。Rov2站位孔隙水溶解Ba2+浓度在上部约440 cm内处于一个较低值（近0 mM），后在约440 cm之下骤然增加，约520 cm处含量达到最大值62.581 μM（图4）。
图 3. 孔隙水阴阳离子深度剖面黑色线为Rov2站位，红色线为PC3站位。Figure 3. Profiles of pore water geochemical parametersThe black line represents the Rov2 core and red line is the PC3 core.3.3 孔隙水中阴离子含量变化
孔隙水中阴离子含量随着深度向下增加（图3），Rov2和PC3两站位孔隙水Cl−含量保持稳定，浓度含量为约560 mM，与现代海水值一致；整体上与Na+、K+具有相似的变化趋势。两岩芯孔隙水SO42−含量分别在上部约245、235 cm保持稳定，接近海水硫酸盐背景值（28.9 mM）；约245、235 cm之下呈线性降低的趋势，约485、410 cm近似降低至0。两站位Br−含量的变化范围不大，整体上略微增加。Rov2和PC3孔隙水PO43−含量具有相似的变化趋势：上层PO43−浓度接近0，随着深度的增加，PO43−浓度逐步增加，在约245、235 cm迅速增加，均在SMI界面附近PO43−含量达到最大值（Rov2岩芯在约615 cm处为19.116 mM；PC3岩芯在约410 cm处为12.326 mM）。
3.4 扩散通量
研究区Rov2和PC3两站位孔隙水的各组分浓度剖面均为非线性，并不适用于Fick第一定律计算其扩散通量。在约300 cm处两站位孔隙水的SO42−、Ca2+、Mg2+、Sr2+、Ba2+浓度以及DIC、δ13CDIC值均发生较大变化（图3）。因此，我们仅计算了两站位300 cm以下一定深度范围内的扩散通量，此时孔隙水各组分浓度呈现相对线性变化，即代表了在相关生物地球化学过程中各组分所产生的瞬时扩散通量[66]。根据公式（3）、（4）计算得到，Rov2站位孔隙水SO42−、Ca2+、Mg2+的扩散通量分别为67.403、11.692、10.806 mol·m−2·ka−1，溶解Ba2+的扩散通量为329.182 mmol·m−2·ka−1。PC3站位孔隙水SO42−、Ca2+、Mg2+的扩散通量分别为94.226、15.120、8.713 mol·m−2·ka−1。
4. 讨论
4.1 硫酸盐还原作用
在海底冷泉系统中，绝大多数甲烷都在硫酸盐甲烷过渡带（SMTZ）内被硫酸盐驱动的甲烷厌氧氧化作用所消耗，甲烷和硫酸盐分别作为电子供体和受体的浓度迅速降低。本研究中，Rov2和PC3沉积物孔隙水SO42−含量分别从245和235 cm开始下降，至485和410 cm处降低至接近0值，此时DIC含量达到最大、δ13CDIC降至最小值，同时Ca2+、Mg2+、Sr2+等离子也分别降至最小值（图2-4）。这说明该层位内发生了强烈的AOM作用，Rov2和PC3柱状沉积物的SMI界面分别为485和410 cm，这与我们通过线性外推拟合计算获得的SMI深度值一致[14, 67-69]。
除了AOM之外，OSR作用也能够消耗硫酸盐，我们利用反应式（1）、（2）中OSR与AOM两个作用生成HCO3−与消耗SO42−化学计量数之比（简写为RC:S）恒为2∶1、1∶1的特点，并进行DIC自生碳酸盐校正[14, 70-71]，对冷泉活动区AOM的相对贡献进行了区分[5, 12, 14-16, 22, 70-71]。图5中，在Rov2浅表层0～235 cm，SO42−接近于海水值，RC:S大于2∶1；在245～485 cm段RC:S介于2∶1与1∶1，AOM作用的参与改变了孔隙水的RC:S之比；380 cm以下，RC:S落在AOM作用1∶1直线上，AOM作用占据主导地位。PC3站位0～235 cm的数据点主要分布在2∶1左右，SO42−主要被OSR消耗而减少；235 cm向下RC∶S在1∶1左右，以AOM消耗SO42−为主。
图 5. Rov2和PC3岩芯中经碳酸盐沉淀校正的溶解无机碳生成量（ΔDIC+ΔCa2++ΔMg2+）与硫酸盐消耗量（ΔSO42−）的关系图（简写为RC:S）以典型海水值（DIC为2.1 mM，Ca2+为10.3 mM，Mg2+为53.2 mM，SO42−为28.9 mM）为参考[12]，计算产生的溶解无机碳量或消耗的硫酸盐量。对角线分别表示1∶1的AOM和2∶1的OSR。Figure 5. Plot of dissolved inorganic carbon produced corrected for carbonate precipitation (ΔDIC+ΔCa2++ΔMg2+) versus sulfate consumed (ΔSO42−) in Rov2 and PC3 cores (abbreviated as RC:S)Typical seawater values (2.1 mM for DIC, 10.3 mM for Ca2+, 53.2 mM for Mg2+, and 28.9 mM for SO42−) were taken as references to calculate the amounts of dissolved inorganic carbon produced or sulfate consumed[12]. Diagonal lines indicate: the RC:S ratio is 1:1 for AOM and 2:1 for OSR, respectively. Gray square represents Rov2 core and red circle is PC3 core.值得注意的是，在PC3下部数据点多落在AOM的1∶1线之下（约0.7∶1），这一结果并不常见，我们认为出现的原因可能是过量硫酸盐被消耗或DIC发生损失。PC3站位稳定的Cl−浓度表明取样沉积物中并不存在水合物分解释放淡水的情况（图3）。Haese等[22]将这一情况归结于研究区高度富Ca2+的盐水流入并向深层迁移从而弥补了沉积物中的Ca2+损失；但本研究区未有证据支持盐水的存在，且PC3岩芯Ca2+剖面连续变化，并不存在外部阳离子的加入。根据附近站位Rov2的实测溶解Ba2+含量（约70 μM）及计算的Ba2+通量（329.18 mmol·m−2·ka−1），较小的变化范围表明重晶石的析出对硫酸根的消耗十分有限。此外，Liu等[71]在采集孔隙水样时加入了HgCl2试剂，H2S与之反应生成H+，酸性环境下也会中和掉孔隙水中的部分DIC。但本次研究中，我们未对水样进行任何处理且PC3的δ13CDIC值变化与Rov2站位相似，表明水样预处理与微生物作用也并非造成该现象的原因。通过对比Rov2，我们发现PC3下部孔隙水DIC含量明显偏低但δ13CDIC值相似（图2），认为可能是由于PC3底部受到某些原因导致过饱和的DIC环境中发生了CO2脱气，这与实际提取岩芯时观察到底部存在两段中空的气层的现象相佐证，当然也存在其他未知反应途径过量消耗DIC的可能，但具体情况仍需要后续的进一步分析。
Rov2和PC3两站位位于非冷泉中心，其SMI深度比琼东南冷泉活动中心的13～20 cm远深得多（表1）[71]，稍浅于神狐海域与东沙海域7～10 m[32, 70]。根据AOM主导时消耗CH4与SO42−的化学计量数之比恒为1∶1（反应式（1）），我们利用向下硫酸盐通量近似估计了向上甲烷通量的大小[17]。Rov2和PC3站位向上甲烷通量分别为67.4、97.2 mol·m−2·ka−1（图6），这与大陆斜坡冷泉渗漏区扩散甲烷通量均值（0.05～800 mol·m−2·ka−1，均值约100 mol·m−2·ka−1[9]）接近，高于神狐海域甲烷的通量值13.8～20mol·m−2·ka−1[32, 70]，远远低于海马冷泉活动中心向上的甲烷通量值 1882.5 、 2110.6 mol·m−2·ka−1[71]。这表明AOM作用进行过程中，两站位SMI深度与甲烷（或硫酸盐）通量之间均表现出高度的负相关性[17-18, 72]。下伏沉积物中富甲烷流体的向上运移，促进AOM消耗大量硫酸盐，SMI随之发生移动深度变浅；且向上甲烷通量高，SMI浅，渗漏活跃；向上甲烷通量低，SMI深度大[17-18, 25, 72]。这充分说明，甲烷通量的大小是控制该区域SMI深浅的关键因素。
表 1. 南海各水合物区SMI深度与甲烷通量Table 1. Comparison of SMI depths and methane fluxes in each site in the South China Sea| Show Table
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4.2 冷泉流体渗漏及碳酸盐岩矿物特征
冷泉流体渗漏发生AOM会伴随着大量DIC的生成，而孔隙水中的DIC及δ13CDIC所携带的巨大信息能够帮助我们进一步追踪、了解孔隙水溶解DIC的碳来源[5, 67-68, 73]。一般来说，海相沉积物孔隙水DIC的主要来源有[73-75]：① 沉积埋藏过程中从海水进入到沉积物的DIC，δ13CDIC值接近0；② 有机质降解产生；③ 甲烷参与的AOM作用。
图2中，PC3站位0～235 cm，孔隙水DIC的δ13CDIC值在−10‰左右略有下降，介于海水δ13C（约为0）与有机质δ13C（南海约为−20‰[76]）之间，孔隙水DIC主要来源于海水和有机质两部分，这与本文4.1节中PC3上部主要为OSR反应生成DIC的结论相一致。Rov2站位上部0～235 cm，孔隙水DIC的δ13CDIC值极度亏损（−40‰～−20‰），且该站位上部的RC:S＞2∶1（本文4.1节），DIC含量相对于PC3站位略高，这表明可能存在额外的DIC加入，且新加入的DIC具有较轻的碳同位素组成。琼东南盆地同时发育有热成因甲烷和生物成因甲烷，其中热成因甲烷所具有的δ13C值为−39.9‰～−34.8‰[77]，生物成因甲烷碳同位素组成通常低于−60‰[78-79]。在Rov2沉积物约485 cm和PC3的410 cm处，δ13CDIC值分别为−48.5‰和−44.5‰（图2），表明该层位孔隙水DIC主要来源于甲烷参与的AOM。
前人用DIC × δ13C与DIC线性回归来确定后期加入孔隙水DIC池的δ13C值[68, 25, 80-81]，直线斜率δ13Cadded即代表了底水自埋藏以来加入到初始DIC池中的δ13C值[81-82]。如图7所示，Rov2顶部在0～235、245～485 cm的δ13Cadded值分别为−55.242‰、−56.805‰，PC3顶部在0～235、235～410 cm的δ13Cadded值分别为−32.029‰、−56.773‰。两个站位0～235 cm的δ13Cadded值明显不同，说明二者具有不同的碳源：Rov2上部后期加入的δ13Cadded极度亏损，这与本文4.2节中新加入的DIC具有较轻碳同位素组成的结论一致。而245～485 cm、235～410 cm处两站位的δ13Cadded相差不大，表明二者孔隙水DIC具有相同的碳源；且新加入的DIC其δ13Cadded值为−58‰～−55‰，证实了上文甲烷参与AOM反应生成的DIC具有较轻的碳同位素组成。
图 7. δ13C×DIC与DIC线性拟合计算得到δ13Caddeda: Rov2岩芯，b: PC3岩芯。Figure 7. δ13Cadde calculated using the linear regression of δ13C×DIC vs. DIC in cores Rov2 (a) and PC3 (b)随着AOM作用的发生，大量DIC生成过程中，促进了孔隙水中的Ca2+、Mg2+等碱土阳离子与之结合，发生碳酸盐矿物沉淀，因此自生碳酸盐沉淀也成为海底渗漏环境的常见产物及海相碳循环的重要碳汇[83-86]。Rov2与PC3沉积物中也存在清晰可辨的碳酸盐颗粒，研究区内富甲烷沉积物中广泛存在自生碳酸盐矿物。利用孔隙水DIC和Ca2+、Mg2+、Sr2+等可以进一步明确碳酸盐的析出过程及矿物学特征[28-29, 32, 87]。
以Rov2站位为例，随着深度的增加，孔隙水Sr2+/Ca2+、Mg2+/Ca2+（图4）逐渐增加；AOM占主导后（约300 cm）变化更加明显；孔隙水中Ca2+的亏损相对于Mg2+变大，意味着自生碳酸盐的沉淀以方解石为主[32, 88]。且在Sr2+/Ca2+与Mg2+/Ca2+图中（图8），Rov2孔隙水数据点落在高镁方解石的组成线附近，沉积物中以析出高镁方解石为主，这与其他冷泉及天然气水合物赋存区有着相似的矿物类型[32, 88]。由于方解石和文石沉淀会受到孔隙水SO42−的抑制作用，且方解石受到的抑制作用要大于文石[89]，通常高镁方解石和白云石会优先在低的SO42−浓度、低渗漏强度的环境中（即SMI及以下深度）发生沉淀，而文石则在渗漏强烈并靠近沉积物表层的具有高CH4通量和高SO42−浓度的环境中析出[21, 29, 70, 90-94]。Rov2站位向上甲烷通量为67.4 mol·m−2·ka−1，与大陆边缘渗漏的平均通量值近似，渗漏活动的强度中等，SMI埋深约485 cm，这也与高镁方解石生成的环境特征相一致。
图 8. Rov2站位孔隙水Sr2+/Ca2+与Mg2+/Ca2+图两条直线表示文石或高镁方解石沉淀过程中孔隙水Sr2+/Ca2+与Mg2+/Ca2+相对于海水成分的变化关系。Figure 8. Plot of pore water Sr2+/Ca2+ vs. Mg2+/Ca2+ ratio of Rov2 coreThe two lines indicate the changes in the Sr2+/Ca2+ to Mg2+/Ca2+ relationship in the pore water with respect to the composition of seawater that occur during precipitation of either aragonite or high Mg-calcite.4.3 AOM对孔隙水Ba2+、Br−、PO43−的影响
近年来，随着测试技术的不断提升，孔隙水Ba微量元素、营养盐组分以及同位素等的研究也逐渐展开，AOM对其影响也逐渐被人们认识[18, 34-36, 38, 66, 70]。
在渗漏活动区，向上释放的流体中一般含有较高的溶解Ba2+[38, 91]，这主要是由于受到AOM的显著影响，沉积物与孔隙水固液相之间Ba元素发生活化、迁移与再沉积：渗漏甲烷向上迁移过程中，消耗大量硫酸根离子，尤其是SMI界面附近，孔隙水硫酸根浓度处于极低值，极大程度上改变了重晶石等含Ba矿物的溶解度，使得沉积物中大量的溶解Ba2+被释放，孔隙水中Ba2+含量发生异常变化[34, 37, 92]。
图4中Rov2站位上部0～400 cm内孔隙水溶解Ba2+处于较低水平，而随着深度向下AOM的发生使得孔隙水极度亏损SO42−（SMI界面处接近于0），沉积物中Ba释放，从而使得Rov2站位沉积物中溶解Ba2+在SMI界面的含量出现急剧增加[34, 92]。我们计算了Rov2站位孔隙水Ba2+通量为329.18 mmol·m−2·ka−1，远高于Hu等[66]在神狐海域的Ba2+通量（47.1 mmol·m−2·ka−1），Rov2站位内下部孔隙水中较高的浓度含量及通量均表明南海海马冷泉区存在沉积物Ba富集的有利条件，至于沉积物中“Ba锋”的存在及具体深度，这都需要对沉积物进行进一步的Ba测试与分析。在Rov2沉积物520 cm以下，孔隙水Ba2+浓度迅速降低，最后稳定在18 μM左右，超出SMTZ后孔隙水Ba2+与固相之间达到的新的局部平衡浓度[35, 91]。
研究表明，甲烷浓度高的环境中，Br−、PO43−等微生物相关元素也会表现出一定的异常[33]。一方面，海水Br−与PO43−来源于有机质降解释放，岩芯上部发生有机质降解反应，较为强烈的生物活动造成孔隙水中Br−、PO43−含量升高。二者的浓度含量一定程度上受到沉积物有机质数量和活性的影响，而这些要素又显著制约着水合物区甲烷的生成，因此可利用二者并结合其他相关指标很好地指示当地有机质富集程度、生烃潜力与微生物活动强度的大小[33, 93]。另一方面，在冷泉活动区，随着深度增加，沉积物中有机质含量逐渐降低，矿化度增加[30]，AOM参与及SMI下部存在的产甲烷作用也可以产生PO43−（反应式为2CH2O(PO43−) → CO2 + CH4 + 2PO43−[15]），这些都使得PO43−的浓度梯度随着深度向下进一步增加。
前人通过对比正常沉积物层孔隙水发现，水合物区的孔隙水中的Br−、PO43−明显偏高，且PO43−浓度梯度曲线存在与SO42−相反的变化趋势，认为其可作为一个新的地球化学指标示踪天然气水合物的存在[33, 41]。研究区Rov2站位上部0～400 cm、PC3站位上部0～300 cm，孔隙水Br−、PO43−随着深度的增加表现出略微增加的趋势（图4），孔隙水PO43−含量高于ODP航次994、997站位水合物区的PO43−含量[94]，图3中二者随深度的变化特征相反，进一步证实研究区孔隙水中Br−、PO43−的分配受到OSR和AOM作用的影响[30, 41]。SMI附近，两站位的PO43−含量出现不同程度的升高，Br−含量也表现为随着深度增加而升高，可能反映了沉积物中有机质含量随着埋深向下而降低，导致其矿化度增加，冷泉活动的发生改变了孔隙水中Br−、PO43−的变化特征[30, 95]。SMI之下，两站位下部孔隙水DIC及δ13C表明岩芯还未到达甲烷生成带，PO43−的变化并非由于甲烷生成作用所引起。不过由于在冷泉渗漏区受到孔隙水中SO42−浓度梯度、甲烷浓度与下伏水合物的平衡和甲烷上升通量间的制约，较低的SO42−水平最终使得PO43−的增加趋于平缓并出现下降[33]。
5. 结论
通过对琼东南活动冷泉区两个重力活塞柱进行孔隙水地球化学分析，Rov2和PC3站位发生了较为明显的甲烷渗漏活动。Rov2站位约245 cm、PC3站位约235 cm开始，孔隙水硫酸盐浓度随着深度向下线性降低，(∆DIC+∆Ca2++∆Mg2+):∆SO42−（RC:S）比值关系表明，随着深度向下，AOM作用在两站位逐步占据主导地位。Rov2和PC3站位的SMI深度分别为约485、410 cm，计算的向上甲烷通量分别为67.4、97.2 mmol·m−2·ka−1，较高的甲烷通量显著制约着SMI深度。受到甲烷渗漏的影响，约245 cm和235 cm之下，AOM产生的DIC为两站位孔隙水DIC的主要来源，且具有较轻的碳同位素组成。依据Sr2+/Ca2+与Mg2+/Ca2+值，Rov2站位自生碳酸盐矿物主要以高镁方解石为主。孔隙水中Ba2+、Br−和PO43−等组分受到特殊生物地球化学作用的影响在SMI附近表现出正异常现象，这与冷泉渗漏活动有着密切的联系。
致谢：感谢中国海洋大学海底科学与探测技术教育部重点实验室张洋、李东永、王楠老师给予的支持与帮助。
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图 5 Rov2和PC3岩芯中经碳酸盐沉淀校正的溶解无机碳生成量（ΔDIC+ΔCa2++ΔMg2+）与硫酸盐消耗量（ΔSO42−）的关系图（简写为RC:S）
Figure 5.
表 1 南海各水合物区SMI深度与甲烷通量
Table 1. Comparison of SMI depths and methane fluxes in each site in the South China Sea
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