Research progress and prospects on the dating of pelagic clay
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摘要:
深海黏土广泛分布在水深超过碳酸盐补偿深度(CCD)以下的大洋盆地中，其沉积速率十分缓慢，只有少量的生物组分(主要是生物磷灰石)被保存，通常具有较高的稀土元素含量；海洋沉积物常用的磁性地层与生物地层相结合的定年手段通常不能有效使用。因此，深海黏土沉积年龄框架的建立一直存在巨大的困难和挑战，严重阻碍了对沉积环境演化和稀土超常富集机制等方面的深入研究。本文回顾总结了20世纪以来逐步发展应用的多种深海黏土定年方法，主要包括磁性地层、鱼牙87Sr/86Sr定年、鱼牙U-Pb定年、10Be测年、230Thex测年、187Os/188Os定年、鱼鳞石生物地层、恒定Co通量模型以及常用的地层对比方法。这些方法各具优缺点，单一使用以上任何一种定年方法几乎都难以获得完整可靠的年龄框架。因此，综合运用多种定年方法，对获得的年龄框架进行系统对比和验证，将会更为有效地提高深海黏土年龄框架的可靠性。
Abstract:Pelagic clay, which is extensively distributed in the ocean basins below the carbonate compensation depth, exhibits slow sedimentation rate and contains only a small amount of preserved biogenic components (primarily biogenic apatite). The commonly used dating methods that combine magnetic stratigraphy with biostratigraphy in marine sediments cannot be effectively applicable. As a result, the establishment of a age framework for pelagic clay has been hindered by enormous difficulties and challenges, which seriously limits the researchers in geoscience to thoroughly investigate the evolution of sedimentation environment and the mechanisms of hyper-enrichment in rare earth elements in pelagic clay. In this article, we reviewed various dating methods for pelagic clay used since the last century, including mainly: magnetostratigraphy, fish teeth 87Sr/86Sr dating, fish teeth U-Pb dating, 10Be dating, 230Thex dating, 187Os/188Os dating, ichthyolith biostratigraphy, constant Co-flux model, and commonly used stratigraphic correlation methods. Each method has own advantages and disadvantages, and it is often difficult to acquire a complete and reliable age framework using any of the above methods alone. Consequently, systematic comparsion and validation for age framework obtained by intergrating multiple dating methods will be more efficient in improving the relability of an age framework for dating pelagic clay.
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Keywords:
- pelagic clay /
- dating method /
- magnetostratigraphy /
- fish teeth/ichthyolith /
- integrated chronology
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天然气水合物是在低温、高压条件下，甲烷等气体与水作用形成的笼形晶体化合物[1]，分布广、储量大、能量密度高，是最为重要的替代能源之一。我国陆地和海洋中的水合物资源相当丰富[2-3]。根据地质构造和储层条件，水合物藏可分为4类（Ⅰ—Ⅳ类）[4-5]。如图1所示，其中，Ⅱ类水合物藏由水合物层和下伏水层组成，顶底为非渗透层。从地质学、地球化学及热力学等角度分析，Ⅱ类水合物藏是分布最为广泛的一种类型，最有希望得到大规模开发利用[6]。
从经济和技术角度看，降压法[8]和热激法[9]是实际水合物藏开采最为可行的方式，常用于现场的开采试验[10-14]。但与常规油气藏不同，水合物藏在开采过程中会发生相变，即固态的水合物吸热分解为可动流体（气和水），单一的降压法或热激法的作用效果往往十分有限。Moridis等[15]研究了定流量抽取地层水，以实现Ⅱ类水合物藏储层降压的方法，但产气效果并不理想。Gao等[16]开展的降压分解实验也表明，在单纯降压条件下，Ⅱ类水合物岩心的生产指标欠佳。杨圣文[17]研究了定井底流压结合井筒加热进行Ⅱ类水合物藏开采的方法，分析了不同井底流压下的开采效果，结果表明，降低井底流压能改善水合物藏的开采效果。Moridis等[18]提出了单井条件下，分阶段进行井筒加热和注热水开采Ⅱ类水合物藏的方法，但施工程序较为复杂，实际应用存在一定困难。Reagan等[19]对Moridis等提出的方法进行了敏感性分析，结果表明孔隙度、储层非均质性等对开采效果有明显影响。
实验研究[20]和数值模拟研究[21]都表明，热水驱替法能够结合降压法和热激法的优点，是水合物藏开采的一种有效方法。对Ⅱ类水合物藏而言，将热水驱替和井组[22]结合具有明显优势。鉴于Ⅱ类水合物藏分布的广泛性，且相关开采模拟仍有待加强，本文结合实际水合物藏参数，使用数值模拟方法研究了热水驱替方式下Ⅱ类水合物藏的开采规律，并与降压法的开采效果进行了对比，以加强对水合物资源开发利用的认识。
1. 研究方法
1.1 开采方法
采用五点井网，热水驱替开采Ⅱ类水合物藏的示意图如图2所示。由于水合物层的渗透率很低、注入性较差，中心注入井在下伏水层的上部区域射孔，向储层注入热水；4口生产井定井底流压生产。在生产过程中，生产井近井地带的低压环境和注入井持续的热水注入使得储层中的水合物大量分解，分解得到的气体运移至生产井采出。
1.2 数值模拟方法
由于水合物藏开采试验不多，数值模拟是目前研究水合物藏开采动态规律的主要手段。HydrateResSim（HRS）是专门用于水合物藏开采的开源学术代码[23]，它考虑了水合物藏开采过程中的相变、传热、多相渗流等机理，能够对降压法、热激法等方式下的水合物藏开采进行有效模拟[24]。在HRS中，相变过程采用主变量变换法处理，数学模型的离散采用积分有限差分法，离散得到的非线性方程组使用Newton-Raphson方法迭代求解。
本文在HRS基础上开展模拟研究，研究的水合物假设为单一的甲烷水合物，体系考虑4相3组分，相包括气相（G）、水相（A）、水合物相（H）和冰相（I），这4相均为储层孔隙中的一部分。其中冰相和水合物相为不可流动相，水相与气相为可流动相，其流动遵循达西定律。组分包括甲烷组分（m）、水组分（w）和水合物组分（由甲烷组分和水组分根据水合数表示，本文水合数取值为6）。其中，水相、气相和水合物相中都存在水组分和甲烷组分，冰相中只存在水组分。
组分κ的质量守恒方程如式（1）所示：
$$\frac{\partial }{{\partial t}}\left(\sum\limits_{\beta = A,G,H,I} {\varphi {S_\beta }{\rho _\beta }X_\beta ^\kappa } \right) + \nabla \cdot \left(\sum\limits_{\beta = A,G} {X_\beta ^\kappa {{\vec F}_\beta }}\right) = {q^\kappa }$$ (1) 式中，φ为储层孔隙度；Sβ为相β（=G, A, I, H）的饱和度；ρβ为相β的密度，kg·m−3；
$X_\beta ^\kappa $ 为组分κ（=m, w）在相β中的质量分数；${\vec F_\beta }$ 为相β的质量流速，kg·m−2·s−1；qκ为组分κ的源汇项，kg·m−3·s−1。体系的能量守恒方程如式（2）所示：
$$\begin{split}&\frac{\partial }{{\partial t}}\left( {(1 - \varphi ){\rho _R}{C_R}T + \sum\limits_{\beta = A,G,H,I} {\varphi {S_\beta }{\rho _\beta }{H_\beta }} + {H_d}} \right) + \nabla \cdot \\ & \left\{ { - \left( {(1 - \varphi ){K_R} + \sum\limits_{\beta = A,G,H,I} {\varphi {S_\beta }{K_\beta })} } \right)\nabla T + \sum\limits_{\beta = A,G} {{H_\beta }{{\vec F}_\beta }} } \right\} = {q^e}\end{split}$$ (2) 式中，ρR为岩石的密度，kg·m−3；CR为岩石的比热容，J·kg−1·K−1；Hβ为相β的焓，J·kg−1；Hd为水合物的形成/分解焓，J·m−3；KR为岩石的导热系数，W·m−1·K−1；Kβ为相β的导热系数，W·m−1·K−1；qe为热量的源汇项，J·m−3·s−1。
模拟研究时，采用平衡模型描述水合物的分解与形成，采用的相对渗透率模型[23]和毛管力模型[23]分别如式（3）、式（4）所示：
$${k_{rA}} = \min \left\{ {{{\left[ {\frac{{{S_A} - {S_{irA}}}}{{1 - {S_{irA}}}}} \right]}^{{n_A}}},1} \right\}$$ $$\begin{aligned}{k_{rG}} = \min \left\{ {{{\left[ {\frac{{{S_G} - {S_{irG}}}}{{1 - {S_{irA}}}}} \right]}^{{n_G}}},1} \right\}\end{aligned}$$ (3) 式中，krA为水相相对渗透率；SA为水相饱和度；SirA为束缚水饱和度；nA为水相相对渗透率递减指数，本文取3.572[25]；krG为气相相对渗透率；SG为气相饱和度；SirG为束缚气饱和度；nG为气相相对渗透率递减指数，本文取3.572[25]。
$${p_c} = - {p_{c0}}{\left[ {{{\left( {\frac{{{S_A} - {S_{irA}}}}{{1 - {S_{irA}}}}} \right)}^{ - 1/\lambda }} - 1} \right]^{1 - \lambda }}$$ (4) 式中，pc为气水间毛管力，Pa；pc0为模型参数，本文取2×103 Pa[25]；λ为模型参数，本文取0.45[25]。
2. 研究实例
2.1 数值模拟模型
参考Mallik地区Ⅱ类水合物藏的参数[26]建立基础模型（如图3所示。模型大小为165 m×165 m×90 m，从上到下，依次由顶部非渗透层、水合物层、下伏水层和底部非渗透层组成。采用直角网格系统，x，y，z方向对应的网格数分别为33×33×22。模型参数如表1所示。采用五点井网热水驱替方式进行模拟研究，生产时间为3年。其中，注入井和生产井的井径均为0.1 m。各生产井的射孔范围为[−1 m，15 m]，以3 MPa进行定井底流压生产；注入井的射孔范围为[−1 m，0 m]，以200 t/d的速度注入50 ℃的热水。
表 1 基础模型参数Table 1. Basic model parameters of the Class Ⅱ hydrate reservoir参数 水合物层 下伏水层 顶底非渗透层 厚度/m 15 15 30 绝对渗透率/10−3 μm2 1 000 1 000 0 孔隙度 0.35 0.35 0 水合物饱和度 0.7 0 − 含水饱和度 0.3 1 − 底部初始压力/MPa 10.67 − − 底部初始温度/℃ 13.3 − − 温度梯度/（℃/100 m） 3.0 束缚水饱和度 0.20 − 束缚气饱和度 0.02 − 由于相变的存在，水合物在开采过程中会与周围环境进行大量的热交换，热效应显著。而水合物藏的顶底非渗透层虽然几乎没有渗透性，但通常赋存着大量热量，在开采过程中能够以热传导的方式为水合物的分解提供能量[27]，因而对水合物藏的开采效果有直接影响。在本文的模型中，顶底非渗透层均为30 m，这个厚度足以刻画开采过程中的热效应[28]。
2.2 开采规律
2.2.1 产气动态
产气速率和累产气、分解气速率和累分解气曲线分别如图4和图5所示。从图4可以看出，产气速率可大致分为两个阶段（阶段①和阶段②），呈现先快速上升，然后以较快速度下降至趋于相对稳定的变化规律。在阶段①，经50 d达到峰值产气速率约为7.0×104 m3/d。在阶段②，产气速率先以较快速度降低，然后逐渐趋于稳定，在生产末期，产气速率约为5×103 m3/d。分解气速率的变化与产气速率变化相同（图5）。对应产气速率和分解气速率的变化，累产气和累分解气前期快速上升，而后近似线性增加。截至生产结束，累产气1.31×107 m3，累分解气1.33×107 m3。除极少量溶于地层水中的气体外，Ⅱ类水合藏中的采出气全部来源于水合物的分解，因此，分解气几乎被全部采出（＞99%），地层中的残余极少。
出现上述变化的原因主要是：在阶段①，由于定压生产以及水层中水的产出，储层压力下降较快，生产井近井地带的水合物快速分解。同时，热水的注入也导致储层水合物的大量分解，但产出的气体主要来自生产井近井地带水合物的分解。在阶段②，随生产时间的增加，在生产井的降压作用和注入井的热水作用（图6）下，水合物持续分解，但由于水合物饱和度持续下降，总体产气速率逐渐降低。
储层温度场的变化如图6所示（剖面图）。可以看出，随生产时间的增加，水合物的分解导致温度降低，模型的温度整体呈逐渐下降趋势，但注入井近井地带由于热水的注入，高温区域不断扩大，能够持续为水合物的分解提供能量。
2.2.2 水合物饱和度
模型水合物饱和度场的变化如图7所示（剖面图）。可以看出，随生产时间的增加，在生产井的降压作用和注入井的热水作用下，生产井周围、水合物层与下伏水层接触面、注入井周围的水合物饱和度逐渐降低；至生产末期，相当一部分水合物已经分解，剩余的水合物主要存在于井间地带。顶部非渗透层携带的热量也促进了水合物的分解，因而水合物层顶部（即顶部非渗透层与水合物层的交界处）的水合物分解明显。当生产结束时，水合物层的平均水合物饱和度降至0.176。
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图 7 水合物饱和度场注：为展示水合物层水合物饱和度的变化，示意图中只显示了部分网格。Figure 7. Hydrate saturation field evolution2.3 开采效果对比
鉴于降压法在水合物藏开采方法中具有重要地位，为对比热水驱替开采水合物藏的效果，建立了降压法开采的对比模型。在降压法开采模型中，将中心注入井关闭，生产井以定井底流压方式生产（3 MPa），其他储层参数及生产参数不变。
在热水驱替和降压开采条件下，水合物藏的气体采出程度、水合物分解程度及累积气水比如图8所示。其中，气体采出程度为采出气量与气体总储量的比值，如式（5）所示：
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图 8 气体采出程度、水合物分解程度和累积气水比Figure 8. Gas recovery percent, hydrate dissociation percent and gas-water ratio$$\eta = \frac{{{N_G}}}{{{R_G}}}$$ (5) 式中，η为气体采出程度；NG为采出的甲烷量，m3；RG为水合物藏中甲烷初始总储量（包括水相中的溶解气以及水合物中蕴藏的甲烷总量），m3。
水合物分解程度为分解的水合物量与水合物总储量的比值，如式（6）所示：
$$\beta = \frac{{{D_H}}}{{{R_H}}}$$ (6) 式中，β为水合物分解程度；DH为分解的水合物量，kg；RH为水合物藏中的水合物总储量，kg。
累积气水比为累积产气量与累积产水量的比值，如式（7）所示：
$$r = \frac{{{V_G}}}{{{m_A}}}$$ (7) 式中，r为累积气水比；VG为累积产气量，m3；mA为累积产水量，t。
从图8可以看出，截至生产结束，相对于降压法，热水驱替的气体采出程度和水合物分解程度更高（分别高出约40%和24%），且热水驱替条件下的气体采出程度和水合物分解程度均大于60%，处于较高水平。两种模型的差异来源于外源热水的注入。相对于降压模型，热水驱替模型中热水的注入扩大了水合物的分解范围，携带的热量提高了水合物的分解程度，相应地气体采出程度也较高；在降压和热水的综合作用下，开采指标优于降压模型。因此，热水驱替对Ⅱ类水合物藏的开采具有一定的适应性，开采指标较好。但热水驱替的产水量较高，其累积气水比仅为降压开采的1/3，这是由外源热水的注入导致的不可避免的结果，因而开采过程中的水处理是十分重要的问题。
3. 结论
（1）热水驱替开采Ⅱ类水合物藏时，水合物藏的产气速率和分解气速率首先快速上升，然后以较快速度下降至趋于相对稳定；累产气和累分解气先快速上升，然后近似线性增加。气体采出程度和水合物分解程度处于较高水平（＞60%）。
（2）热水驱替对Ⅱ类水合物藏的开采具有一定的适应性。相对于降压开采，在热水驱替条件下，水合物藏的气体采出程度更高，水合物饱和度的降低更显著；但累积气水比较低，伴随较大的产水量。
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图 1 研究区域图
a：本文所涉及的部分站位分布图，b：太平洋沉积物类型分布图。海洋沉积物类型数据来自文献[11]。各站位参考信息见表1，其中GC62的数据尚未发表。
Figure 1. The study areas
a: The distribution of some stations covered in this article; b: the distribution of sediment type in the Pacific Ocean Data of marine sediment type is from [11]. The reference information of each station is shown in Table 1, with unpublished data for GC62.
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图 2 西太平洋GC62孔中部分样品的交变退磁正交矢量图和剩磁衰减图
实心方块代表水平面投影图，空心方块代表垂直面投影图。
Figure 2. Orthogonal vector projection and remanence attenuation of alternating demagnetization
Solid squares: horizontal projections; hollow squares: vertical projections
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图 4 40 Ma以来的海水87Sr/86Sr参考曲线（a）和PC01、GC1901孔中鱼牙87Sr/86Sr深度变化（b）
a中数据来自文献[43]，PC01孔的鱼牙87Sr/86Sr数据来自文献[12]，WP41孔的鱼牙釉质87Sr/86Sr数据来自文献[18]，GC1901孔的鱼牙87Sr/86Sr数据来自文献[14]。
Figure 4. The 87Sr/86Sr reference curve of seawate since 40Ma (a) and the 87Sr/86Sr chang with depth of fish teeth in cores PC01 and GC1901 (b)
Data sources: the 87Sr/86Sr reference curve: from reference [43], fish teeth 87Sr/86Sr data of core PC01: from [12], fish teeth enamel 87Sr/86Sr data of core WP41: from [18]; fish teeth 87Sr/86Sr data of core GC1901: from [14].
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图 5 鱼牙釉质U-Pb定年结果
a：保存完整的鱼牙化石照片，b：WP41孔鱼牙釉质U-Pb定年结果，c：Site 1218孔中的鱼牙釉质U-Pb定年结果，d：Site 1218孔中的鱼牙U-Pb年龄同古地磁年龄和有孔虫年龄之间的对比。a，b, c, d均改绘自文献[18]。
Figure 5. Dating results of fish teeth enamel
a: Photo of well-preserved fish teeth fossil, b: U-Pb dating results of fish teeth enamel in WP41, c: U-Pb dating results of fish teeth enamel in Site 1218, d: comparison of U-Pb ages between fish teeth in Site 1218 and paleomagnetic / foraminiferal ages.a,b, c, and d are redrawn from [18].
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图 6 北太平洋GPC3孔年龄综合
a：古地磁年龄、10Be年龄、鱼鳞石生物地层、Co模型年龄及K—E年龄点汇总，各年龄数据来源见表1；b：古地磁年龄与10Be年龄比较。
Figure 6. Integrated age in core GPC3 in the North Pacific
a: Paleomagnetic age, 10Be age, ichthyolith biostratigraphy, Co model age, and K-E age. Data sources are shown in Table1; b: comparison between paleomagnetic ages and 10Be ages.
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图 7 西太平洋core C孔和GC18孔年龄模型综合
a: 西太平洋core C孔年龄模型图，改绘自文献[16]；b: 西太平洋GC18孔11～15.5 Ma年龄模型和沉积速率图；c: GC18孔50点平滑的Ba元素含量随深度变化图；d: 11～16 Ma的轨道偏心率变化图（蓝线）和经调谐后的GC18孔Ba元素随年龄变化图（红线）。图b、c、d均改绘自文献[17]。
Figure 7. Integrated age model in cores C and GC18 in the Western Pacific
a: Age model plot of core C in the Western Pacific, redrawn from [16]; b: 11～15.5 Ma age model and sedimentation rate plot of GC18 in Western Pacific; c: plot of 50 point smoothed Ba element content in GC18 varies with depth; d: plot of orbital eccentricity variation during 11～16 Ma (blue line) and plot of smoothed Ba element changes with age (red line). b, c, d are redrawn from [17].
表 1 部分深海黏土沉积物定年研究总结
Table 1 Summary of some pelagic clay dating studies
区域和站位 位置 定年方法 年龄范围和或深度 平均沉积速率
/(mm/ka)备注 参考文献 北太平洋
GPC330.33˚N、
157.82˚W鱼鳞石生物地层 上新世—古新世 0.2～0.3 误差约±1～5 Ma [5] 磁性地层 布容极性期
更新世2.2
1.7[7] 鱼牙87Sr/86Sr 用以上两种方法来进行验证 [8] 10Be 0～6 m
6～10 m约1.2
约0.5以1.387 Ma为半衰期重新
进行计算[6] 恒定Co通量模型 晚古新世—中中新世 0.2 显著低于更新世时期的沉积速率 [3] 铱元素异常 异常高值指示K-Pg边界(约65 Ma) [3] 北太平洋
PC0132.5˚N、
141.2˚W鱼牙87Sr/86Sr 0～10.7 m (0～24 Ma) 0.45 误差±1～3 Ma [12] 东赤道太平洋
PC078.8˚N、
135.4˚W鱼牙87Sr/86Sr 0～15 m (0～19 Ma)
0～4 m (深海黏土)
4～16 m (硅质黏土)
0.3
2.0误差<±2 Ma [13] 东赤道太平洋
GC19019.78˚N、
154.97˚W鱼牙87Sr/86Sr 21～32 Ma 2.3 误差约±0.8～3 Ma [14] 西太平洋
WPPC1902-0818.29˚N、
149.84˚E磁性地层 0～6 m (0～2.59 Ma) 2.3 棕黄色深海黏土 [15] 西太平洋
core C20.22˚N、
161.48˚E230Thex
自生10Be/9Be
磁性地层0～3.1 m (0～11.6 Ma)
0～1.2 Ma
1.2～11.6 Ma
1.67
0.125多种测年方法获得的沉积
速率一致[16] 西太平洋
GC1816.90˚N、
162.18˚E自生10Be/9Be
磁性地层
轨道调谐1.8～5.4 m
(11～15.4 Ma)0.1～2.5 [17] 西太平洋
WP4123°N、
158°E鱼牙U-Pb定年 2.2～6.5 Ma 1.4 误差±1～2 Ma [18] 西太平洋
PC1122.98˚N、
154.02˚E187Os/188Os
磁性地层9～12 m 0.43～1.02 假设187Os/188Os识别的E2-E3边界位于磁性地层内某一极性时期 [19-20] 南太平洋
U136522.85˚S、
165.65˚W磁性地层
Co通量模型0～6 m
10～18 m约1
＜0.2[21-22] 
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