10.1 Introduction

Traditional stable isotope geochemistry usually refers to isotopes of C, H, O, N and S. With the development of multicollector inductively coupled plasma‒mass spectrometry (MC–ICP–MS), high-precision measurements of more isotopes have led to the thriving of so-called non-traditional stable isotope geochemistry (Halliday et al. 1995). The distinctive geochemical features of these non-traditional stable isotopes, such as diverse geochemical activity, various concentrations in different geological reservoirs, redox sensitivity, and the links to biological activities, make them unique tracers for different geological processes (Teng et al. 2017).

The formation of authigenic minerals at seeps involves element migration and corresponding isotope fractionation, such as Fe and Mo for pyrite and Mg and Ca for carbonate. In addition, the geochemical processes in the cold seep system leads to changes in the redox state in the environment and involves the isotope fractionation of redox-sensitive elements, such as Mo. The seep system is an excellent “laboratory” for studying the behaviors of these isotopes in the natural environment. These isotopic compositions of different seep deposits have been successively analyzed in recent years. The South China Sea is one of the most investigated seep systems, and various non-traditional stable isotope geochemical techniques have been used to explore various geochemical processes. Iron isotopic compositions of anaerobic oxidation of methane (AOM)-derived pyrite from several sites provide constraints on the relationship with AOM activity and pyrite formation (Lin et al. 2017, 2022). The Mo enrichment and isotopic composition of sediments and carbonates provide new insights into the geochemical cycling of Mo (Lin et al. 2021). Magnesium isotopic compositions of pore water from “Haima” and seep carbonate from the Dongsha area put forward a new understanding of the behavior of Mg isotopes during carbonate precipitation (Jin et al. 2021, 2022). Studies of the Ca isotope geochemistry of dolomite from the Dongsha area have constrained the process of carbonate formation (Wang et al. 2012, 2013).

10.2 Fe Isotopes

Iron is a major widely distributed element in the silicate Earth. It has three oxidation states, namely, metallic iron, ferrous iron and ferric iron, which lead to its complex chemical and isotopic behaviors. Iron has four stable isotopes, with mass numbers of 54, 56, 57 and 58, which represent 5.845%, 91.754%, 2.119%, and 0.282% of the total mass, respectively (Meija et al. 2016). The Fe isotope ratios are usually reported as δ56Fe:

$$\updelta {}^{56}\mathrm{Fe}=\left[\frac{{\left({}^{56}\mathrm{Fe}/{}^{54}\mathrm{Fe}\right)}_{\mathrm{sample}}}{{\left({}^{56}\mathrm{Fe}/{}^{54}\mathrm{Fe}\right)}_{\mathrm{standard}}}-1\right]\times 1000$$

where the standard usually refers to IRMM-14. Another standard, the average of terrestrial igneous rocks (IgRx), was previously used to define the isotopic composition of Fe (Beard and Johnson 1999). The conversion of δ56Fe values relative to the two standards can be adjusted according to Beard et al. (2003a):

$${\updelta {}^{56}\mathrm{Fe}}_{\mathrm{IRMM}-14}={\updelta {}^{56}\mathrm{Fe}}_{\mathrm{IgRx}}+0.09\mathrm{\permil }$$

in the following text, δ56Fe values refer to δ56FeIRMM-14 unless noted otherwise.

Pyrite is a common mineral in anoxic marine sediments, and its sulfur isotopic composition serves as an archive for constraining the biochemical cycle of sulfur and the evolution of ocean chemistry (Canfield and Teske 1996; Farquhar et al. 2000; Bekker et al. 2004; Johnston et al. 2006; Pasquier et al. 2017). The geochemical cycles of sulfur and iron are intimately coupled and recorded in sedimentary pyrite (Rouxel et al. 2005; Johnson et al. 2008; Hofmann et al. 2009; Liu et al. 2020). Pyrite in seep systems can present as framboids, overgrowths and euhedral crystals in sediments or enclosed in seep carbonates (e.g., Peckmann et al. 2001; Lin et al. 2016). The variable sulfur isotopic compositions of pyrite are closely related to AOM activity (e.g., Jørgensen et al. 2004; Borowski et al. 2013). The study of Fe isotope geochemistry may be useful to better understand the formation process of pyrite in seep systems. The Fe isotopic composition of pyrite depends on (1) the Fe isotopic composition of Fe sources and (2) the fractionation of Fe isotopes during the formation of pyrite (Butler et al. 2005; Guilbaud et al. 2011; Dauphas et al. 2017). Terrigenous active iron minerals are the main iron source for pyrite formation in continental margin sediments (Berner 1970; Poulton et al. 2004). Both theoretical calculations and experiments show that ferrous iron is 56Fe-depleted compared to ferric iron that is either dissolved or contained in minerals (Johnson et al. 2002; Welch et al. 2003). Aqueous Fe2+ is released into pore water during dissimilatory iron reduction (DIR), with fractionation of −0.5‰ to −2.95‰ between aqueous Fe2+ and Fe oxides (Dauphas et al. 2017 and references therein). In addition, abiotic reductive dissolution by sulfide and iron-mediated anaerobic oxidation of methane (Fe-AOM) can also release isotopically lighter Fe into pore water (Riedinger et al. 2014; Egger et al. 2015). Hydrogen sulfide produced by organoclastic sulfate reduction and AOM reacts with dissolved iron or reactive iron minerals to form metastable intermediates, such as mackinawite, pyrrhotite and greigite, which are finally converted into pyrite (Jørgensen and Kasten 2006). The ∆56FeFe(II)-FeS are −0.52 ± 0.16‰ in equilibrium at 2 °C and +0.85 ± 0.3‰ in kinetic fractionation (Butler et al. 2005; Wu et al. 2012). The fractionation of Fe between FeS and pyrite is +2.20 ± 0.70‰ (Guilbaud et al. 2011). Combining all these processes, potential fractionation between aqueous Fe2+ and pyrite can vary from −3.1‰ to +0.5‰ depending on different extents of pyritization (Guilbaud et al. 2011).

To date, the reported δ56Fe values of bulk pyrite from seeps in the South China Sea range from −1.48‰ to +0.38‰ (n = 42) (Lin et al. 2017, 2018), similar to the range of variation in δ56Fe values of other sedimentary pyrites (Fig. 10.1). The Fe isotopic compositions of bulk sediment at seeps from the South China Sea (δ56Fe =  +0.038‰ to +0.122‰, n = 17; Fig. 10.1) are also close to those of other modern marine sediments (δ56Fe =  − 0.12‰ to +0.23‰; Beard et al. 2003b; Severmann et al. 2006; Fehr et al. 2010). However, in situ analyses of the Fe isotopic compositions of pyrite aggregates of different types from seeps by LA–MC–ICP–MS revealed large variations in δ56Fe values ranging from −1.09‰ to +1.90‰ (n = 160) (Lin et al. 2018).

Fig. 10.1
A dot plot compares delta superscript 56 F e to various areas. Pyrite of L A-M C-I C P M S has the highest values in the South China Sea at around 2.0. The rest of the areas plot both pyrite and bulk sediment. The Baltic Sea has the lowest value for the pyrite at around negative 1.7.

Comparison of iron isotopic compositions of bulk sediment and pyrite from modern marine environments. Data from the South China Sea are from Lin et al. (2017, 2018). Data for the Peru upwelling area are from Scholz et al. (2014). Data from the Baltic Sea are from Fehr et al. (2008, 2010). Data from the Black Sea and Santa Barbara are from Severmann et al. (2006). Note that the pyrites from the South China Sea were hand-picked, while those from other studies were obtained through a chemical leaching procedure; see details in Severmann et al. (2006)

The iron pools in marine sediments include iron oxyhydroxides (Feox), magnetite (Femag), carbonate (Fecarb), silicate (Fesil), and pyrite (Fepy). Highly reactive iron (FeHR) is defined as the sum of Feox, Femag, Fecarb and Fepy, and Fepy/FeHR is introduced to indicate the extent of pyritization (e.g., Scholz et al. 2014; Lin et al. 2017, 2018). Researchers have suggested that iron oxyhydroxides and magnetite are the major iron sources for the precipitation of pyrite at seeps in the South China Sea (Lin et al. 2017, 2018). The enrichment of 56Fe in pyrite is consistent with the increased degree of pyritization in the South China Sea (Fig. 10.2a). It has been suggested that δ56Fe values of pyrite are controlled by both sulfide availability and iron pools in seep systems (Lin et al. 2018). A positive correlation between δ56Fe and δ34S values of pyrite has been observed, and pyrite affected by AOM and organoclastic sulfate reduction can be distinguished by different Fe and S isotopic compositions (Fig. 10.2b; Lin et al. 2017). In situ analysis of δ56Fe indicates that the later-stage overgrowths and euhedral pyrites are more enriched in 56Fe than the earlier formed framboids (Lin et al. 2018). This finding has been explained by Rayleigh-type distillation owing to the relatively Fe-limited environment within sediment during the formation of later-stage pyrite (Lin et al. 2022). The combination of Fe and S isotopic compositions of pyrite from seeps serves as a potential proxy to trace the different biogeochemical processes in early diagenesis.

Fig. 10.2
A scatter and cluster plots of delta superscript 56 F e versus f e subscript p y by F e subscript h r. a. The plots are distributed in an increasing trend indicated by the diagonal arrow. b. 2 clusters for O S R and A C M affected. The A C M affected plot has the highest value at around (55, 0.35).

Enrichment of 56Fe in pyrite with increasing extent of pyritization (a) and plot of δ56Fe against δ34S of pyrite from the Shenhu area in the South China Sea (b). Modified after Lin et al. (2017, 2018)

10.3 Mo Isotopes

Molybdenum is a transition element with a wide range of redox states (–IV to +VIII), leading to its high degree of chemical reactivity. Molybdenum (IV) and Mo (VI) are most common on Earth's surface. Molybdenum has seven stable isotopes (mass numbers of 92, 94, 95, 96, 97, 98 and 100) with similar abundances of approximately 10% to 25% (Hlohowskyj et al. 2021). The isotopic composition of Mo is expressed as δ98Mo relative to NIST-SRM-3134 = 0.25‰ (Nägler et al. 2014):

$$\updelta {}^{98}\mathrm{Mo}=\left[\frac{{\left({}^{98}\mathrm{Mo}/{}^{95}\mathrm{Mo}\right)}_{\mathrm{sample}}}{{\left({}^{98}\mathrm{Mo}/{}^{95}\mathrm{Mo}\right)}_{\mathrm{NIST}-\mathrm{SRM}-3134}}-1\right]\times 1000+0.25\mathrm{\permil }$$

where +0.25‰ accounts for the offset from the in-house standards used previously (Nägler et al. 2014).

Molybdenum has a long residence time of ~0.44 to 0.8 Ma and is generally ubiquitous (~105 nM) in oxygenated oceans (Miller et al. 2011; Nakagawa et al. 2012; Hlohowskyj et al. 2021). The Mo isotopic composition (~2.3‰) is also homogenous in modern open seawater (Hlohowskyj et al. 2021). Molybdenum is redox sensitive and has characteristic isotopic fractionation (Kendall et al. 2017). The Mo isotopic signatures in sediments have been used as a proxy to reconstruct marine redox conditions (e.g., Poulson et al. 2006; Scott and Lyons 2012; Kendall et al. 2017 and references therein). In oxic environments, Mo mainly presents as molybdate (MoO42−) in seawater (Siebert et al. 2003). The dissolved Mo can easily absorb to metal oxides, leading to ∆98Moseawater-solid of approximately 0.8‰ to 3.2‰ (e.g., Barling et al. 2001; Siebert et al. 2003; Wasylenki et al. 2008; Scholz et al. 2017). Under anoxic/sulfidic conditions, MoO42− transforms to thiomolybdates (MoOxS(4-x)2−) facilitated by hydrogen sulfide, which is easily captured by iron sulfide or sulfur-rich organic matter and then stored in sediments (Helz et al. 1996, 2011; Tribovillard et al. 2006). The removal of Mo from aqueous environments is mainly controlled by the availability of dissolved hydrogen sulfide (Kendall et al. 2017). At [H2S]aq > 11 μM, Mo is quantitatively transformed into tetrathiomolybdate and scavenges into sediments (Erickson and Helz 2000). The sediments record the Mo isotopic composition of ambient water (Nägler et al. 2011). Under conditions with [H2S]aq < 11 μM, sediment yields a large variation in Mo isotopic composition due to incomplete sulfurization from molybdate to tetrathiomolybdate (Erickson and Helz 2000; Neubert et al. 2008; Nägler et al. 2011).

The AOM in seep systems create strongly sulfidic environmental conditions in pore water, with [H2S]aq easily exceeding 11 μM in the case of high methane flux (e.g., Gieskes et al. 2005, 2011; Joye et al. 2010). This facilitates the accumulation of authigenic Mo in sediments or seep carbonates. The enrichment of Mo in seep sediment and the covariation of Mo and U in seep carbonate are strongly associated with redox states and the intensity of AOM (Hu et al. 2014; Chen et al. 2016). Seeps may serve as an important sink in the marine geochemical cycle of Mo (Hu et al. 2015). Therefore, Mo isotope geochemistry can provide new insights into tracing the sedimentary environment in seep systems. To date, only sparse data on the Mo isotopic composition of sediments and carbonates at seeps have been reported (Lin et al. 2021). The δ98Mo values of bulk sediment and carbonate nodules from the South China Sea range from +0.2‰ to +2.0‰ (n = 36) and +1.4‰ to  +2.1‰ (n = 5), respectively (Fig. 10.3; Lin et al. 2021). The large variation in Mo isotopic compositions of sediment was ascribed to the mixing of authigenic Mo sequestered in the course of AOM and detrital Mo in sediment unaffected by seepage. In addition, the later-formed pyrite may also cause additional Mo accumulation and influence the δ98Mo of bulk sediment (Lin et al. 2021). For seep carbonate with a narrow range of δ98Mo values, the offset between δ98Mocarbonate and δ98Moseawater is close to that of strongly euxinic conditions (Lin et al. 2021). Considering the much higher H2S concentration in active seepage than the critical Mo speciation threshold of ~11 μM, δ98Mocarbonate was thought to have potential for tracing the Mo isotopic composition of seawater (Lin et al. 2021).

Fig. 10.3
A graph of modern seawater versus delta superscript 98 M O. It plots carbonate, skeletal and nonskeletal, oxygenated, mildly oxygenated to anoxic, weekly euxinic, strongly euxinic, seep carbonates, and seep sediments in the South China Sea. The seep sediment has the highest range, from negative 1 to 3.

Molybdenum isotopic compositions of seep carbonates and sediments from the South China Sea (Lin et al. 2021) and other sinks of Mo in the modern oceans (modified after Kendall et al. 2017)

10.4 Mg Isotopes

Magnesium is widely distributed in the silicate Earth, hydrosphere and biosphere. It has three stable isotopes, i.e., 24 Mg, 25 Mg and 26 Mg, with typical abundances of 78.97%, 10.01% and 11.03%, respectively (Meija et al. 2016). The isotopic composition of Mg is usually reported as δ26Mg relative to the DSM3 standard (Young and Galy 2004):

$$\updelta {}^{26}\mathrm{Mg}=\left[\frac{{\left({}^{26}\mathrm{Mg}/{}^{24}\mathrm{Mg}\right)}_{\mathrm{sample}}}{{\left({}^{26}\mathrm{Mg}/{}^{24}\mathrm{Mg}\right)}_{\mathrm{DSM}3}}-1\right]\times 1000$$

Magnesium is the second most abundant cation in seawater, with a long residence time of ~13 Ma (Li 1982). The Mg isotopic composition of modern seawater is −0.83 ± 0.09‰ (Ling et al. 2011). Carbonate precipitation represents a sink of Mg in the ocean and is an important part of the oceanic Mg cycle (Higgins and Schrag 2015). The Mg isotopic composition of marine carbonates serves as a tool for unraveling the geochemical cycling of Mg in the ocean on geological timescales (e.g., Tipper et al. 2006; Higgins and Schrag 2012, 2015; Fantle and Higgins 2014). Carbonate is generally enriched in 24 Mg relative to seawater and displays a large variation in δ26Mg (Teng 2017). Fractionation of Mg isotopes during carbonate precipitation represents a temperature dependence of only approximately 0.01‰ °C−1, while it exhibits a strong mineralogical control (Saenger and Wang 2014). The order of enrichment from light to heavy isotopes is calcite < magnesite < dolomite < aragonite (Saenger and Wang 2014). In addition, the Mg isotopic composition of carbonate could also be influenced by other factors, such as the precipitation rate, inorganic and organic ligands in solution, and biological effects (Saenger and Wang 2014 and references therein; Schott et al. 2016; Mavromatis et al. 2017).

Most seep carbonates from the South China Sea are composed of microcrystalline high-Mg calcite (HMC) and aragonite cement. In addition, a significant amount of microcrystalline dolomite occurs at some seep sites (Feng et al. 2018 and references therein). The content of MgCO3 in some HMCs can even reach 38 mol%, occurring in association with minor amounts of dolomite (Han et al. 2008). Previous studies have confirmed that the presence of hydrogen sulfide provided by AOM can prominently affect the incorporation of Mg into carbonate minerals (e.g., Naehr et al. 2007; Zhang et al. 2012; Lu et al. 2018, 2021; Tong et al. 2019). Therefore, seep deposits are good subjects for understanding the behavior of Mg isotopes during carbonate precipitation in natural environments. The Mg2+ in seawater diffusing into pore water is the main source of Mg in seep carbonates. The δ26Mg of pore water from an active seep in the South China Sea ranges from −0.88‰ to −0.71‰ (from 0 to 8 meter below the seafloor), similar to that of seawater (Jin et al. 2022). A slight increase in the pore water δ26Mg values with depth indicates the precipitation of authigenic carbonate. This occurs because the sufficient replenishment of Mg from seawater is far greater than the consumption of Mg in pore water during precipitation of authigenic carbonate (Jin et al. 2022). Lu et al. (2017) and Jin et al. (2021) analyzed the Mg isotopic composition of seep carbonates from the South China Sea. A comparison of the δ26Mg values of seep carbonates and other marine carbonates is presented in Fig. 10.4. The δ26Mg values of seep carbonates display a smaller range of variation relative to those of other marine carbonates. Additionally, the biologically mediated vital effects on the Mg isotopic composition are not obvious compared to biodetrital carbonates based on current data. Carbonates in Lu et al. (2017) consist mainly of dolomite, with relatively consistent Mg isotopic compositions (δ26Mg =  −2.45‰ to −2.15‰, n = 11). However, the δ26Mg values of HMC in pipe-like seep carbonates vary from −3.42‰ to − 2.63‰ (n = 10), and a decreasing trend of δ26Mg values from the periphery to the inner portion of the “pipe” was observed (Jin et al. 2021).

Fig. 10.4
A graph of delta superscript 26 Mg. The y axis is split into 3 parts. The top portion for biogenic carbonate includes high and low Mg calcite, and aragonite. The middle portion for abiogenic carbonate includes island and modern marine dolomite, and bulk marine carbonate. The last potion for seep carbonate is split into dolomite and high mg calcite.

Comparison of magnesium isotopic compositions of carbonates. Data on abiogenic carbonate are from Saenger and Wang (2014) and references therein. Data on island dolomites are from Hu et al. (2022). Data on modern marine dolomites are from Higgins and Schrag (2010), Fantle and Higgins (2014), Mavromatis et al. (2014) and Blättler et al. (2015). Data on dolomite and HMC in seep carbonate are from Lu et al. (2017) and Jin et al. (2021), respectively

The correlation of inorganic δ13C and δ26Mg values of seep carbonates is considered to be related to the dissolved hydrogen sulfide in pore water produced by AOM (Lu et al. 2017; Jin et al. 2021). The existence of dissolved H2S has been confirmed to promote the incorporation of Mg into carbonate in precipitation experiments (Zhang et al. 2012). Hydrogen sulfide derived from AOM in seep systems has also been deemed to facilitate the formation of carbonate minerals with high Mg content, especially dolomite (Lu et al. 2015; Tong et al. 2019). The observed different trends of δ13C and δ26Mg values refer to opposite trends in the degree of Mg isotope fractionation with enhanced intensity of AOM (Fig. 10.5). The dissolved H2S may affect the Mg isotopic composition of carbonates, although the mechanism is still not well understood. This may provide new insights into the application of carbonate Mg isotopic compositions during geological time of widespread oceanic anoxia.

Fig. 10.5
A graph plots delta superscript 26 m g versus delta superscript 13 C. Two arrows pointing from the bottom right to the top left are titled enhanced A O M and one arrow pointing from the top right to the bottom is titled enhanced A O M. Plots for Shenhu and S W Taiwan lie along the former, while plots for Dongsha lie along the latter.

Correlation of inorganic δ13C and δ26Mg values of seep carbonates from the South China Sea (modified after Jin et al. 2021). Data from the Shenhu and SW Taiwan areas are from Lu et al. (2017). Data from the Dongsha area are from Jin et al. (2021)

10.5 Ca Isotopes

Calcium is abundant in both terrestrial and marine systems. It has five stable isotopes (40Ca, 42Ca, 43Ca, 44Ca and 46Ca) and a radioactive isotope (48Ca) with an extremely long half-life of 1.9 × 1019 years (Griffith and Fantle 2021). The abundance of isotopes from 40 to 48Ca is 96.941%, 0.647%, 0.135%, 2.086%, 0.004% and 0.187%, respectively (Meija et al. 2016). The Ca isotope ratios are expressed as δ44Ca:

$$\updelta {}^{44}\mathrm{Ca}=\left[\frac{{\left({}^{44}\mathrm{Ca}/{}^{40}\mathrm{Ca}\right)}_{\mathrm{sample}}}{{\left({}^{44}\mathrm{Ca}/{}^{40}\mathrm{Ca}\right)}_{\mathrm{standard}}}-1\right]\times 1000$$

Multiple standards are in use to define δ44Ca, such as SRM 915a, SRM 915b, seawater, and bulk silicate Earth (Griffith and Fantle 2021). Data are reported relative to seawater in the following text. The conversion of δ44Ca relative to SRM 915a and seawater is according to Hippler et al. (2003):

$${\updelta {}^{44}\mathrm{Ca}}_{\mathrm{SRM }915\mathrm{a}}={\updelta {}^{44}\mathrm{Ca}}_{\mathrm{seawater}}+1.88\mathrm{\permil }$$

The formation of carbonate plays an important role in the cycling of Ca. The calcium isotopic composition of carbonate has been used to quantify rates of carbonate diagenesis and trace global Ca cycling (Griffith and Fantle 2021). The δ44Ca values of modern marine carbonates generally display isotope fractionations between − 1.8‰ and − 0.8‰ from seawater (Blättler et al. 2012; Fantle and Tipper 2014). Similar to Mg isotopes, carbonate mineralogy and precipitation rate are two dominant factors controlling the calcium isotopic compositions of carbonate (Gussone et al. 2005; Tang et al. 2008; Blättler et al. 2012). The high alkalinity of pore water in seep systems facilitates the formation of authigenic carbonate with a fast precipitation rate (Aloisi et al. 2000; Naehr et al. 2007). The calcium isotopic compositions of seep carbonate may have the potential for understanding the formation process of carbonate during early diagenesis. The diffusion of seawater-dissolved Ca2+ into pore water is the main source of authigenic carbonate at seeps. Carbonate depleted in 44Ca indicates that the precipitation of seep carbonate prefers light Ca isotopes (Fig. 10.6). Mg-calcite in seep carbonate yields higher δ44Ca values than aragonite (Fig. 10.6). The reported δ44Ca values of dolomite-dominated seep carbonate from the South China Sea (−0.67‰ to − 0.41‰, n = 11) are similar to those of Mg-calcite (Wang et al. 2012, 2013; Thiagarajan et al. 2020; Blättler et al. 2021). Wang et al. (2012) suggested that the variation in δ44Ca values can be attribute to supersaturation, precipitation rate and the degree of Ca consumption in semi-closed systems at seeps. The Rayleigh effect on Ca isotope fractionation during the precipitation of seep carbonate can also be indicated by the downward-increasing δ44Ca values of pore water from the Storfjordrenna Trough in Blättler et al. (2021).

Fig. 10.6
A graph plots delta superscript 44 C a. The y-axis is into the South China Sea, Storfjordrenna Trough, the North Sea, Barents, and the Cascadia margin. The South China Sea plots dolomite while the Cascadia margin plots aragonite, and the Storfjordrenna Trough plots mg-calcite.

Comparison of Ca isotopic compositions of seep carbonate. Data from the South China Sea are from Wang et al. (2012, 2013). Data from the Storfjordrenna Trough and the North Sea are from Thiagarajan et al. (2020). Data from the Cascadia margin are from Teichert et al. (2005)

10.6 Summary and Future Studies

In summary, the special and complex sedimentary environments in seep systems could impact the behaviors of isotopes, such as Fe, Mo, Mg and Ca. The iron isotopic compositions of pyrite are strongly related to its formation process and ambient environmental conditions. Isotopic signatures of Mo in seep carbonates and sediments are relevant to Mo migration and transformation with a changing redox state. Magnesium and Ca isotopic compositions of seep carbonate reflect the precipitation of authigenic carbonate, and the hydrogen sulfide derived from AOM may significantly influence the fractionation of Mg isotopes. Non-traditional isotope geochemistry in seep systems is still in its infancy. Whether the special environment at seeps affects the fractionation mechanism of these isotopes is unclear. The application potential of these non-traditional isotopic compositions of seep deposits also remains to be developed. Below are some problems in the field of non-traditional isotope geochemistry in seep systems that could be further studied in the future:

  1. 1.

    Pyrite in seep systems is the result of continuous accumulation along with the dynamic activity of methane seepage. The Fe isotopic compositions of pyrite may be an assemblage of complex biogeochemical processes, especially when multiple seepage events occur. More research is needed to further reveal the information hidden in the Fe isotopic compositions of pyrite.

  2. 2.

    Molybdenum in seep deposits exists in carbonate minerals, sulfide phases, detritus and organic materials. The Mo isotopic signature of bulk seep carbonate is a mixture of the δ98Mo of all these Mo-bearing phases. This causes uncertainty in the application of seep carbonates to trace the Mo isotopic compositions of seawater. Analyzing the Mo isotopic compositions of different phases in seep carbonates by sequential chemical extraction may help to better understand the behavior of Mo isotopes in seep systems.

  3. 3.

    Dissolved hydrogen sulfide plays an important role in the behavior of Mg. Therefore, the AOM in seep systems has been thought to have a possible effect on the fractionation of Mg isotopes during the formation of seep carbonates. The behavior of Mg isotopes during carbonate precipitation under sulfidic conditions could be better constrained by precipitation experiments.

  4. 4.

    Seep carbonates form under a fast precipitation rate. The precipitation rate is an important factor controlling the fractionation of Ca isotopes during carbonate precipitation. However, the Rayleigh effect also obviously influences the δ44Ca of seep carbonates. Therefore, more research is needed to better constrain the precipitation rate of seep carbonate to quantify the effect of the precipitation rate on its Ca isotopic composition.