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自20世纪80年代以来海洋Fe的生物地球化学及其与全球碳循环、海洋生态的关系一直受到广泛关注,并在近些年取得了一系列突破性认识[1-2]。Fe作为生化反应中协助电子传递的关键元素,参与了浮游生物的光合作用、呼吸作用以及固氮作用[2-3]。尽管地壳中Fe是常量元素之一,现代海水中Fe3+及其无机络合离子的溶解度却极低(~0.01 nmol/kg[4]),溶解Fe总浓度在开阔大洋的表层海水中一般不超过0.2~0.5 nmol/kg[5-6],因而成为海洋生物生产的限制性因子之一[7-9]。全球现代大洋超过20%的面积属于富营养低叶绿素海区[7],可能与表层海水“Fe限制”密切相关[7-9]。海洋观测显示“Fe施肥”(Iron fertilization)可以促进固氮和初级生产率的提高[10-11]。著名的铁假说认为,南大洋在冰期时风尘输入Fe的增加可能有效地增强了生物泵,促进了大气CO2浓度的降低[12],并得到了模拟研究的支持[13]。除了对生物活动产生直接影响外,海洋溶解Fe最终会进入沉积物中。这些活性Fe的通量与赋存状态会进一步影响有机质、磷以及一些微量金属在沉积物中的保存[14-15],从而在长时间尺度上调节全球C循环。因此,准确认识大洋溶解Fe的来源是可靠预测不同时间尺度中Fe循环对全球变化反馈的基础。
Fe主要以3种方式输入开阔大洋,包括大陆边缘沉积物孔隙水扩散[16-19]、风尘溶解[20-23]、以及热液喷流[24-30]。除自然活动释放的Fe外,人类活动产生的含Fe气溶胶沉降也是现代海洋Fe的来源之一[31],可能对南大洋部分区域以及大陆边缘海的Fe收支有着重要影响[32]。本文仅限于讨论自然来源Fe的物源通量及其示踪。由于河口的凝絮作用,一般认为河流的溶解Fe很难直接输出到大洋[33]。尽管海水Fe浓度已经发表了约两万个数据点,由于其浓度控制因素非常复杂,且分布高度不均一,至少受到外源输入、生物利用、配体络合、颗粒吸附-解析、洋流混合的综合控制[1-2, 34-35],至今没有数值模型能很好地重现大洋溶解Fe浓度[36]。根据近些年大洋Fe生物地球化学循环的研究进展,本文将简要探讨3个方面的问题:(1)现代大洋溶解Fe来源的通量估计;(2)Fe同位素对物源的示踪;(3)第四纪冰期-间冰期旋回中Fe的物源演变及其争议。
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海水Fe的赋存状态按物理上分为溶解态(定义为可透过0.2或0.45 μm滤膜的Fe)和颗粒态(图1)。从化学形态上看,溶解Fe实际上包含了自由或无机络合离子、有机小分子螯合物,有机大分子螯合物,与微生物如细菌相关的微粒吸附态,生物排泄物微粒吸附态、氧化物微粒、碎屑微粒[2, 35]。其中那些颗粒尺寸小于20 nm,主要与有机小分子(如Fe载体)或者无机阴离子络合的Fe被称为可溶Fe,而更大粒径则为胶态Fe。海水中的溶解Fe大约一半以胶体形式存在,且表层海水中的胶态Fe比例变化范围比深水中的大[35]。无机胶态Fe(如碎屑Fe,氧化物Fe)可能是部分表层海溶解Fe的重要赋存形态,且具一定化学惰性,不易与其他组分发生同位素交换[37-38]。此外,在热液喷口附近,无机胶态Fe也占较大比重[39]。除了这些区域,海水中从可溶Fe到胶态Fe几乎完全被有机配体所络合[34],这些Fe通常活性较强,可以自由交换[37],易被生物利用。溶解Fe一般处于不饱和状态,即水柱中有机配体浓度几乎总是超过溶解Fe的浓度,称为“配体过剩”[34, 40]。 配体过剩的现象在上部海洋特别是Fe含量低而生产力高的海域尤其显著[41]。值得注意的是不同有机配体在海水中的存留时间可能具有较大的变化范围[42]。Volker 和 Tagliabue[42]推测主要产生于海洋表层的小分子有机配体(如Fe载体)可能在几年内就被分解,而深水中较弱的有机配体的存留时间则较长,甚至能达到上千年。
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当风尘沉降进入表层海洋后,无定形的Fe氢氧化物等活动性最强的Fe会发生部分溶解并被生物利用。风尘Fe的溶解度受控于光化学氧化还原过程[43]以及颗粒Fe化学组分的影响[44]。传统上认为风尘是输送到北太平洋、东赤道太平洋、南大洋表层海洋Fe的主要途径,有利于缓解这些地区表层的Fe限制,促进生物生产[21]。风尘向大洋输入的溶解Fe通量(F风尘)可以通过公式(1)表示:
${F_{\text{风尘}}} = \sum\limits_i^n {{\theta _i}{\varPsi _i}} $
(1) 其中n代表风尘源区的数量,Ψ为每个风尘源区输入海洋的颗粒Fe通量。θ为风尘颗粒Fe在海水中的溶解度。Tagliabue等[36]总结了13个不同的海洋Fe循环模型,发现由于不同模型的θ取值不一,得到的F风尘变化范围很大(1.4~32.7 Gmol/a)。在开阔海域,中低纬大西洋比其他地区Fe含量显著提高(图2),可能代表了撒哈拉风尘的重要贡献[22]。风尘Fe在其他地区的贡献还存在较大争议,比如Winckler等[45]和Tagliabue等[46]分别分析了赤道东太平洋和南大洋的表层海水的Fe收支,认为洋流上涌带来的Fe比风尘输入大得多。
图2显示靠近大陆架的表层海水的Fe含量显著超过1 nmol/kg(如北冰洋、地中海、印尼海区、阿拉伯海以及洋岛附近区域),部分地区甚至超过5 nmol/kg。陆架海水Fe的浓度可能受沿岸海水中有机配体的浓度控制[47],而从陆架沉积物扩散到海水中的Fe通量取决于沉积物水界面附近孔隙水的Fe浓度梯度和溶解Fe的扩散系数[48]。这两个参数与底水氧气含量、沉积物活动性Fe含量、生物灌溉度以及有机质氧化速率等过程密切相关[16, 19]。Dale等[16]基于底层海水O2含量和沉积物有机碳含量提出了新的近似公式,沉积物溶解Fe通量(F沉积物)的积分形式可以表示为:
${F_{\text{沉积物}}} = \mathop{{\int\!\!\!\!\!\int}\mkern-21 mu \bigcirc}\nolimits_s \gamma \cdot \tanh \left( {\frac{{{C_{\rm{ox}}}}}{{{O_{\rm{2BW}}}}}} \right) \cdot ds$
(2) 其中,Cox是有机质氧化速率(mmol/(m2·d−1)),O2BW是底层海水的氧气含量(μM),而γ则是与沉积物中活动性Fe含量相关的参数,该公式忽略了底水氧气含量低时生物灌溉的抑制效应。通常而言,底水氧气含量较低而生物生产率高的地区如东赤道太平洋的最低含氧带,陆架向大洋输出的Fe相对较高,达到每年每平方米数百毫摩尔[49]。Dale等[16]估计的Fe通过沉积物-水界面的总通量为109±55 Gmol/a,其中72 Gmol/a由浅部陆架贡献(<200 m),37 Gmol/a由深部陆架-陆坡(200~2000 m)贡献。近年来Cai等[50-51]发展了224Ra/228Th不平衡法来估算孔隙水中溶质通过沉积物-水界面通量的技术,其最大优势是在不对沉积环境产生扰动的情况下定量获取生物灌溉和分子扩散等过程的贡献。基于这一技术,Shi等[19]发现滨海沉积物中Fe在沉积物-水界面的输出潜力要比传统底栖箱式模型或者单纯分子扩散估计高至少一个数量级,表明陆架作为Fe的输出源区可能比以前的认识更为重要。除了陆架沉积物扩散外,高纬地区的冰川融水或冰筏沉积物的溶解也是海洋溶解Fe的一大来源[53-54]。Zhang等[55]通过对冰川河流的Fe浓度与同位素研究发现,尽管河流本身Fe浓度可以达到数百nmol/kg,但在河口稳定下来进入海洋的溶解Fe浓度可能仅有几个nM,因此,过去研究对冰川融水释放Fe的潜力可能估计过高。目前,模拟研究[36]中所估计的沉积物溶解Fe通量范围为0.6~194 Gmol/a,并通过调节海水中Fe的移除速率使模拟和观测的海水Fe浓度接近一致,但是这些模型缺乏对通量不确定性的评估。
海底热液活动主要发生在洋中脊和弧后盆地。Fe是热液中最富集的金属之一,其浓度可比海水高6~7个数量级[56-57]。Fe一旦从热液喷涌至深水中,会被迅速氧化形成颗粒物,而Fe2+的半衰期可能仅有几分钟至几个小时[58-59],因此传统上认为热液活动对大洋深水溶解Fe的贡献很有限。引言中已述及,越来越多的研究表明热液可能是一个被忽略的重要源区。一方面,热液区微生物活动可以输出大量的有机配体[60],从而络合并稳定Fe[26, 61]。另一方面,一部分热液Fe可以氧化形成纳米粒径的颗粒从而构成溶解Fe的一部分[24, 39]。实际上热液Fe的信号在离喷口区数千千米的地方仍可以被识别出来[27-28]。热液溶解Fe的通量(F热液)可以简单用公式(3)表示:
${F_{\text{热液}}} = \sum\limits_i^n {{\eta _i}{\varTheta _i}} $
(3) 其中,n表示所有热液活动区的数量,Θ表示热液Fe的通量,η表示能转化为溶解Fe的比例。Tagliabue等[30]假设η为0.2%,计算出的热液溶解Fe通量为0.9 Gmol/a。也有部分研究[24, 26]基于对单个热液喷流区的估计,认为热液Fe中溶解Fe的比例可达百分之几。Resing等[28]通过溶解Fe与热液3He通量的关系,估算出输出到深海的热液Fe通量可达4 Gmol/a。此外,Saito等[29]发现南大西洋慢扩张的洋中脊dFe/3He比已报道的快速扩张洋中脊高达80倍,因此,认为热液输出的溶解Fe通量可能要比前人的估计高很多。
从以上讨论不难发现,不同来源的Fe输出通量无论是相对比例还是实际贡献均存在相当大的不确定性,且不同研究差异很大。需要注意的是除了风尘溶解出来的Fe可以参与整个水柱循环外,热液羽状喷流的胶粒Fe和通过沉积物水界面扩散的还原性Fe可能难以远距离传输,因而增加了新的不确定性。
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Fe同位素为识别大洋溶解Fe的来源提供了新的手段。Fe同位素一般用δ56Fe表示:δ56Fe=((56Fe/54Fe)样品/(56Fe/54Fe)IRMM-014−1)×1 000,其中IRMM-014为国际标准。评估不同端元对深海的贡献需要知道不同来源的Fe同位素组成特征以及在海水中传输时发生的分馏。国内文献[62]近期已就全球大洋中Fe同位素分布和潜在物源的同位素特征进行了详细介绍,本文主要从物源示踪方面作一些简要分析。
风尘的活动性Fe同位素特征以及溶解进入海水时发生的分馏还缺乏系统研究,定量估计表层海水的风尘贡献较为困难。一般来说,Fe的氢氧化物在有机配体作用下发生络合溶解时,轻的Fe同位素会被优先溶解释放,反映了动力学分馏过程[63-64]。Wiederhold等[64]报道了针铁矿被草酸淋滤的初始几分钟,Δ56Fe淋滤液-针铁矿有高达−2‰的动力学分馏。但在同位素交换平衡时,有机配体相对无机离子一般倾向于结合重的Fe同位素。Dideriksen等[65]标定了25 ℃下溶液中Fe3+-去铁胺的平衡分馏系数,发现去铁胺络合的Fe比溶液自由Fe的δ56Fe高0.60±0.15‰。在海洋环境中,大西洋中低纬表层海水中溶解δ56Fe可超过(+0.6‰[66]),如果风尘δ56Fe值本身与硅酸盐平均值类似(~0.1‰[67]),则风尘部分溶解中的平衡分馏过程可能占据重要地位。
陆架孔隙水中的Fe在输出沉积物水界面时一般经历了氧化-还原过程,具有显著偏轻的同位素组成[18, 66, 68-70]。这种偏轻的同位素特征容易在大洋中识别出来,例如来自于北大西洋GA03剖面和东太平洋的GP16剖面(图3)。由南美陆架输出深海的溶解δ56Fe可以低至−0.4‰以下,且具有相对高的溶解Fe浓度。类似地,低δ56Fe信号在大西洋两侧靠近陆架的水体中也很明显。水体中溶解δ56Fe与硅酸盐类似的同位素特征则被认为是来自于氧化物非还原性溶解(即不经历Fe3+还原而直接被溶解)[66, 71-72]。
热液贡献的溶解Fe同位素组成并不均一。如图3所示,北大西洋洋脊热液流体δ56Fe可以显著小于−1‰,而东太平洋洋脊热液区δ56Fe仅略低于0‰。热液贡献的溶解Fe的同位素组成可能取决于热液喷出前或在喷口附近初始沉淀的含Fe矿物类型[26, 73-75]。一般硫化物具有相对轻的Fe同位素组成,而氧化物Fe同位素则偏重,使得不同区域热液溶解Fe同位素有较大变化。目前还不清楚同一热液区在长时间尺度上是否会发生Fe同位素演化,而这些信息有可能为地质历史上热液的化学环境提供新的约束。
利用不同源区Fe同位素组成特征的溯源混合模型已成功用于甄别大洋溶解Fe的来源,其中最典型的例子来自于Conway 和 John[66]对北大西洋GA03剖面海水Fe来源的定量解析。他们首先将这一海水剖面分成3个子区域,并认为每个子区域分别主要受到两种不同物源的控制。假设来自非还原溶解的δ56Fe为+0.09‰,热液δ56Fe为−1.35‰,还原性溶解δ56Fe为−1.35‰,风尘溶解δ56Fe为+0.68‰,对每个子区域利用两端源混合模型定量计算了4种来源Fe的贡献。其结果表明撒哈拉风尘对该区海水溶解Fe可能有着71%~78%的贡献,占据主导;而非还原性溶解Fe大约占10%~19%,还原性溶解Fe大约占1%~4%,热液Fe大约占2%~6%。
值得注意的是溶解Fe从海水中移除时可能同样伴随同位素分馏,表现在生物利用Fe[52, 76]、Fe3+的吸附沉淀[77]等过程。大洋胶态组分Fe占总溶解Fe大约一半,胶体聚合沉淀可能也是溶解Fe从海水中移除的重要途径[78],不过这种方式目前还缺少可靠的表征[36]。在Fe循环模型中一般假设深水Fe的移除速率与局域颗粒通量以及溶解Fe的浓度成正比,较少考虑胶体移除机制。海水溶解Fe进入颗粒相时可能伴随显著的Fe同位素分馏[77, 79]。在东太平洋的San Pedro盆地,沉淀与溶解态间的同位素分馏约为−0.8‰,暗示溶解Fe以无机氢氧化物的形式移除是氧化性水体中发生的主要过程[77]。类似地,在开阔大洋的自生Fe氧化物沉淀中,其Fe同位素可能也主要受Fe离子和有机配体之间的同位素分馏所控制。前人通过比较已发表的铁锰结壳表层样品和当地海水的δ56Fe发现,结壳表层和海水存在较为一致的Fe同位素分馏,即Δ56Fe结壳-海水=−(0.77±0.06)‰[80]。因此,利用Fe同位素示踪物源时必须考虑溶解Fe移除过程中的同位素分馏效应。此外,溶解Fe吸附进入颗粒相下沉的过程中,可能存在可逆吸附行为,即表层的Fe同位素信号可以被有效地带入深海中[39, 68]。更可靠地表征Fe的移除机制及其同位素分馏是利用Fe同位素示踪大洋Fe循环的基础,是未来研究的一项重要任务。
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第四纪以轨道尺度的北半球冰期-间冰期旋回为主要气候特征。一般认为,大气CO2浓度在第四纪气候变化中起到了关键性的调节作用,而C与气候之间的相互反馈是古气候研究的核心问题之一。海洋作为表生系统中最大的C库,其生物泵在过去的演化可能显著影响了大气CO2浓度。了解生物生产限制性营养元素Fe在第四纪海洋中的收支,是理解生物泵演化的重要一环。第四纪也是地质历史上沉积记录保存最好且材料最为丰富的时期,这为认识海洋Fe循环在不同时间尺度上对物源变化的响应提供了重要条件。在第四纪气候旋回中,有多种过程可以影响源区向大洋释放Fe的过程。如冰川体积变化引起海平面升降会影响陆架暴露面积,并进一步影响Fe向开阔海的传输;深水的氧化-还原状态可能会伴随着洋流格局的重组出现周期性波动,从而调节Fe从海水水柱中的移除效率以及沉积物中Fe的保存与释放;冰期-间冰期旋回风尘通量或深水上涌强度的改变也会影响表层海水的Fe收支。一些研究还发现热液活动可能也有明显的冰期-间冰期波动[81-83],一般认为是因为海平面周期性变化引起洋中脊压力发生改变,进一步影响岩浆熔融和热异常[84]。
南大洋与赤道东太平洋是深水与大气交换气体的重要区域。伴随着次表层海水上涌,大量的营养被带到表层,但由于Fe含量极低,限制了这些区域的生物生产和营养利用。图4对比了南大洋钻孔冰期-间冰期的沉积碎屑Fe通量、生物生产率指标及大气CO2浓度记录。在冰期气候中南美、南非、澳大利亚等干旱区释放的风尘通量显著增加,使得冰期时亚南极海铁限制减弱,生物泵效率加强,有利于C存储到深海中,引起大气CO2浓度降低[85-86],这与经典“铁假说”的认识一致[12]。类似地,Murray等[87]通过比较赤道西太平洋钻孔沉积总Fe通量和生物硅质沉积通量在轨道尺度上的演化,支持风尘输入缓解了赤道太平洋Fe限制,增加了生物生产率。此外, Loveley 等[88]利用232Th代表风尘通量,用过剩Ba代表生物生产率,重建了赤道东太平洋末次间冰期以来千年尺度上这些指标的演化,表明千年尺度上风尘的“铁施肥”效应同样成立。但是这些认识受到了来自现代海洋研究[46]和古海洋记录[45, 89]的挑战。如Tagliabue等[46]提出热液活动Fe通量大,可能会有效地缓冲风尘输入的变化对南大洋表层Fe收支的影响。热液活动释放的Fe激发南大洋表层生物生产的直接观测在最近也有报道[90]。从地质记录的角度,Winckler等[45]发现过去生物生产率增高的同时伴随着钻孔沉积物Fe/Th的快速升高,认为这反映了洋流而不是风尘带来的Fe。Costa等[89]系统地研究了赤道太平洋在末次冰期和全新世的生物生产率、营养利用率以及风尘的通量变化,认为风尘变化对生物生产率的贡献实际上很有限。
图 4 南大洋钻孔PS75/059-2(太平洋侧)[85]和ODP-1090(大西洋侧)[86]碎屑通量、生物生产率以及大气CO2浓度在最近50万年的变化[94]
Figure 4. Evolution of the detrital flux, biological productivity, and atmospheric CO2 concentrations[94] over the last 500 ka from the Southern Ocean sediment cores PS75/059-2 (Pacific sector)[85] and ODP-1090 (Atlantic sector)[86]
与现代过程研究相比,第四纪大洋溶解Fe的来源研究主要集中在风尘通量对生物生产率的影响方面,而陆架或热液Fe输出与全球环境交互关系的研究较少。Scholz等[91]利用氧化还原敏感元素指标(U、Mo通量、N同位素)分析了赤道东太平洋陆架在末次间冰期以来的古海洋环境与Fe向深海输出通量的关系,发现在上部海洋缺氧过程中Fe向开阔海输出反而更少了。他们认为这与水体缺氧导致表层沉积物形成硫化环境,从而使得Fe高效地与硫结合相关,并推测促进陆架Fe释放仅存在一个很窄的底水氧气含量区间,这与现代观测认识一致[19]。
利用Fe同位素示踪大洋Fe物源演化的研究则更为有限。目前仅有少量以铁锰结壳为载体的论文发表。铁锰结壳生长在大洋海山上,其主要成分之一为直接沉淀的溶解态Fe,因而有潜力保存海水的Fe同位素信息。Zhu等[92]利用铁锰结壳Fe同位素开展了大洋Fe的物源研究,并与Pb同位素时间序列对比发现陆源物质输入是控制北大西洋深水溶解Fe的主要来源。在太平洋,Chu等[93]通过研究日本岛和赤道附近的几块铁锰结壳的Fe同位素在10 Ma以来的演化,认为热液输入对研究区溶解Fe有重要贡献。Horner等[80]利用结壳重建了76 Ma以来北太平洋的Fe同位素演化,发现了δ56Fe的大幅度波动,认为其反映了热液和陆架沉积物溶解的控制。不过,结壳的铁锰氧化物在沉积后发生矿物相转变是否会影响古海水Fe同位素重建的准确性,尚需开展进一步研究。
由于海水Fe来自于不同物源所含的活动性组分的溶解,研究大洋沉积物活动性Fe的同位素演变是定量理解深水Fe收支的另一途径。东太平洋GEOTRACES GP16剖面的颗粒活动性Fe同位素(可被草酸-EDTA提取的组分)表明,在靠近南美陆架的地方,颗粒活动性铁δ56Fe显著偏低,而洋盆中央δ56Fe则更可能受到了热液活动的影响[79]。东太平洋热液向深海贡献的溶解Fe以无机胶粒形式为主,具有较低的同位素交换反应活性,似乎在深海传输过程中保持着稳定的同位素特征[39]。同时,温跃层/中层水中铁的循环很快(一般数年至数十年),生物利用较少且颗粒通量比表层小,可能是大陆边缘溶解Fe转运的快速通道[72, 91, 95-96]。当颗粒物吸附这些Fe沉积于海底后,颗粒表面铁氢氧化物会向热力学更加稳定的结构转化,这将使其活性降低几个数量级[97]。如果颗粒物沉积后没有经历孔隙水氧化-还原反应驱动的活性Fe迁移,我们可以认为沉积物中的活性Fe同位素组成可以反映沉积时的组成。根据IODP的钻孔资料,在开阔大洋中有广泛区域至少在第四纪晚期没有出现显著的孔隙水缺氧,为重建过去颗粒物活动性Fe同位素提供了基础。深海碎屑沉积物δ56Fe接近黄土组成[67],但是活动性Fe组分的同位素组成特征目前并不明确。前人利用钻孔沉积物化学淋滤已初步开展了不同物相的活动Fe同位素分析[98-99]。由于弱淋滤过程本身可能导致同位素分馏[99],可靠地提取和测定不同物相Fe同位素组成还需继续探索。
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在国际痕量元素及其同位素海洋生物地球化学循环研究计划(GEOTRACES)的支持下,近些年海水溶解Fe的生物地球化学研究取得了一系列突破性认识,重构了我们对大洋溶解Fe地球化学循环认识的概念模型(如图5所示)。其主要进展是获得了海水溶解Fe的同位素组成分布,区分了不同来源的Fe同位素组成特征,并认识到有机配体对大洋Fe循环的关键控制作用。在物源研究方面,热液输入对海水Fe的影响得到了越来越多的关注,甚至可能是南大洋深水Fe的主要来源。
尽管如此,陆架-热液-风尘这些源区对大洋溶解Fe的贡献还存在很大的不确定性,而对这些源区的Fe输出如何响应气候变化的理解亦很有限。在地质尺度上,不同学者对于冰期-间冰期和千年尺度风尘输入与“铁施肥”效应的关系存在很大争论,是当下研究的热点。一些研究认为热液活动在冰期、间冰期也可能存在系统变化。如果这些变化叠加了深海氧气含量的变化,那么热液对Fe循环的影响有可能进一步放大。此外,已有的对陆架沉积物水界面Fe扩散通量的各种估计均无法回答这些Fe能在多大程度上到达开阔海。事实上,陆架沉积物孔隙水溶解Fe含量可达μmol/kg级别,而陆架海水本身的溶解Fe浓度仅仅在nmol/kg量级。可以预期的是,绝大多数从沉积物中扩散出来的Fe会很快被氧化并以颗粒形式回到沉积物-水界面。由于开阔海水的溶解Fe几乎都是与有机配体络合而成的稳定形式,陆架水体的有机配体循环或许才是决定沉积物扩散出来的Fe参与大洋循环的限制性和决定性因素。最后,末次冰期以来的气候变化已经有较好的约束,若利用沉积物的活动性Fe同位素和Fe通量综合约束开阔大洋Fe的物源演变,有可能为理解陆架-热液-风尘输出Fe的机制提供可靠答案,这方面的工作亟需开展。
The sources of dissolved iron in the global ocean and isotopic tracing
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摘要: 铁(Fe)作为海洋初级生产所必需的微量和限制性营养元素影响着海洋生物群落结构、生态功能以及碳循环,理解溶解Fe的物质来源及其对气候变化的响应具有重要的科学意义。早期研究多强调风尘输入是维持大洋Fe循环的主要机制。近年来,随着海水Fe分析数据的积累,尤其是痕量元素及其同位素海洋生物地球化学循环研究计划(GEOTRACES)的开展,陆架沉积物和热液活动所释放Fe的贡献开始越来越受到重视。尽管如此,不同物源对开阔大洋溶解Fe的影响依然存在相当的不确定性。以海水溶解Fe的化学组分为出发点,强调有机配体对大洋Fe循环的决定性作用,综述了不同来源Fe的通量估计和第四纪大洋Fe来源的研究争议。铁同位素为理解大洋Fe的物源演变提供了新的工具。讨论了不同物源的Fe同位素特征,并提出结合沉积物的活动性Fe同位素和组分研究可能为理解过去陆架-热液活动-风尘输出与输运Fe的机制提供全新视角。Abstract: Iron is an essential trace element to oceanic primary productivity, which may influence the structure of marine biological community, ecological function, and carbon cycle. It is therefore of great importance to understand the sources and supply of dissolved Fe to the ocean and its responses to the global climate change. Early studies often emphasize dust input as the mechanism to maintain oceanic Fe cycling. In recent years, with the increase in Fe data, especially along with the launch of GEOTRACES program, the important role of dissolved Fe released from continental shelf sediments and hydrothermal activities has been highlighted. Nevertheless, there still remain considerable uncertainties regarding the contribution of Fe to the open ocean from different sources. Our review begun with characterizing the chemical speciation of dissolved Fe, especially of organic ligand in oceanic Fe cycling, and then presented flux estimation of different Fe sources as well as the debates regarding the oceanic Fe fertilization during the time of Quaternary. Iron isotopes provide a new tool for studying the evolution of Fe sources. We have discussed the Fe isotope signatures of different sources, and proposed that the combination of isotopes and speciation analysis of sedimentary reactive Fe might provide a new perspective in understanding the mechanism of Fe export and transport from continental shelf, hydrothermal activity, and dust in the past.
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Key words:
- dissolved Fe /
- sources /
- Fe isotopes /
- hydrothermal /
- dust
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图 1 海水Fe赋存分类示意图
按物理尺寸可以分为溶解Fe和胶态Fe以及颗粒Fe;其中溶解Fe的化学赋存形态如右侧所示。注意图中不同化学形态的粒径区间仅为示意,不完全代表实际尺寸(据文献[2, 34, 35]整理)
Figure 1. Illustration of iron classification in seawater
Dissolved Fe can be classified according to particle sizes as soluble Fe, colloidal Fe, and particulate Fe; the chemical forms of Fe are shown on the right. Note that grain-size ranges of different chemical forms are only for illustrative purposes, not representing actual size (adapted from [2, 34, 35])
图 4 南大洋钻孔PS75/059-2(太平洋侧)[85]和ODP-1090(大西洋侧)[86]碎屑通量、生物生产率以及大气CO2浓度在最近50万年的变化[94]
阴影所示的大气CO2浓度低值时期对应南大洋高的风尘输入通量与生物生产率
Figure 4. Evolution of the detrital flux, biological productivity, and atmospheric CO2 concentrations[94] over the last 500 ka from the Southern Ocean sediment cores PS75/059-2 (Pacific sector)[85] and ODP-1090 (Atlantic sector)[86]
图 5 以大西洋-南大洋剖面为例的海水Fe生物地球化学循环与物源示意图(据文献[1]修改)
南大洋的Fe限制和中低纬的N、P限制是由于高纬富营养、缺Fe的深水上涌,而中低纬风尘通量高但N、P耗尽引起的。南大洋生产的过剩有机配体可能随着水体传输到中低纬和深水,从而远程调节全球大洋的Fe循环
Figure 5. Illustration of the biogeochemical cycle of Fe in the ocean with the Atlantic-Southern Ocean profile as an example (modified from [1])
The Fe limitation in the Southern Ocean and N-P limitation in the mid-low latitudes are due to the upwelling of nutrient-rich, Fe-depleted deep waters at high latitudes while N-P are depleted with high dust input in the mid-low latitudes. Excessive ligands are produced in the Southern Ocean which can be transported to the mid-low latitudes and deep waters, and in turn will remotely regulate the global Fe cycle
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