Sensitivity, reliability, and implications of SSM measurements
Although the magnetic field produced by the studied speleothem thin section was quite weak (0.1 nT or less), our SSM measurements successfully provided information on distributions of variable magnetization at a resolution of ~ 100 µm. The new procedure of drift corrections applied in this study is based on taking the average along with line scan measurements between predefined threshold values (excluding local magnetic anomalies produced by contaminated dust particles, etc.) away from the median value in the marginal band area used for drift correction.
SSM magnetic images of NRM indicated that NRM before AF demagnetization (Fig. 3a, b) is dominated by negative anomalies (magnetized in the downward direction relative to the surface of the thin section), which is the opposite polarity to the other SSM magnetic images (i.e., 5mT, 10mT, and 30mT; Fig. 3c–h). This suggests that secondary magnetization was acquired after speleothem formation, which was removed by AF demagnetization at 5 mT. This is confirmed by the magnetization measurements on the thin section with the SRM (Additional file 5: Figure S2). The results of the SSM magnetic imaging and SRM measurements suggest that the averaged coercivity of the speleothem is ~ 10–20 mT, and the NRM can be reduced to ~ 20% by AFD at 30mT. Meanwhile, the previous studies on speleothems from Canada, the UK, and Mexico (Latham et al. 1979, 1986, 1987) and from Japan (Morinaga et al. 1989) reported the removal of viscous remanent magnetization (VRM) by AF demagnetization using peak fields of 5–10 mT. Speleothems from China showed the removal of VRM by AF peak field of 30 mT (Openshaw et al. 1997). Thus, the secondary magnetization could be adequately removed in the speleothem by AF demagnetization.
SSM magnetic images of IRM suggest that the concentration of magnetic minerals is higher in the gray layer than in the white part (Fig. 2h, i). We estimated magnetic concentration for the grayish layer and the white part based on the estimated magnetic moments (Fig. 2i). As mentioned in the methods section, the thickness of the samples is 200 µm, and we calculated a magnetic moment on a 200 µm grid. Thus, the saturation magnetization of the grayish layer is 2.54 × 10–1 A/m, and that of the white part is 6.23 × 10–3 A/m. As described below, the main magnetic remanence carriers in the speleothem studied could be magnetite. Assuming that all magnetic remanence carriers in the speleothem are single-domain magnetite (Ms: 446 kA/m; Mrs/Ms = 0.5), we can estimate that the magnetic concentrations per volume in the speleothem of the gray layers and white part as 1.14 × 10–6 and 2.79 × 10–8, respectively.
The above-described experiments demonstrated that SSM is capable of imaging very weak magnetic field differences induced from differences in laminated structures of a speleothem at a resolution of ~ 100 µm.
Magnetic minerals in the speleothem
Magnetic mapping of the speleothem from the stalagmite of Anahulu cave with the submillimeter scale resolution with SSM (Fig. 2a–c) clearly showed a difference in the magnetization distributions among the grayish surface layer, the white inner part, and the reddish-brown layer (Fig. 2e–f). A series of rock magnetic results suggest that the main magnetic remanence carriers in the speleothem studied are magnetite and maghemite, and the minor magnetic remanence carrier is hematite (or ε-Fe2O3; see discussion below). Similar results have been obtained in other stalagmites (e.g., Strauss et al. 2013; Font et al. 2014; Jaqueto et al. 2016).
The cooling curves of ZF-cycling for AAC-B3 and W3, and the warming curve of ZF-cycling and FC/ZFC curves for AAC-W3 show the characteristic inflection points at ~ 100 K (Fig. 4a, b, d, and e), which can be attributed to the Verwey transition (125 K for stoichiometric magnetite, e.g., Jackson and Moskowitz 2021), a diagnostic feature of magnetite. The relatively sharp inflection observed for the white inner part (AAC-W3) suggests that their main magnetic carrier is relatively stoichiometric magnetite. In contrast, the vague shift to lower temperatures or absence of the Verwey transition for the grayish surface layers (AAC-B) is considered as the result of highly oxidized (deviation from stoichiometry) magnetite (e.g., Jackson and Moskowitz 2021). The difference between the ZF-cycling curves of TRM and ThD for AAC-B3 and M was not distinguished from zero at 300 K (Fig. 4g, h). These features can indicate the absence of goethite (Lascu and Feinberg 2011).
On the other hand, the inflection point at ca. 200 K in the TRM curve for AAC-M seems to be related to a Morin transition, which can suggest the presence of hematite (Fig. 4f, i). In general, Morin transition occurs at ca. 260 K, but it can be shifted toward lower temperatures for impure hematite (e.g., de Boer et al. 2001; Liu et al. 2010). Thus, the inflection point at ca. 200 K in TRM, which can be observed only in AAC-M, may suggest that hematite exists in the reddish-brown layer. Meanwhile, the inflection point can alternatively be explained by the presence of ε-Fe2O3. The magnetic property of ε-Fe2O3 corresponds to a high coercive field and low unblocking temperature (e.g., Tuček et al. 2010; López-Sánchez et al. 2017; Kosterov et al. 2020b). Furthermore, the occurrence of ε-Fe2O3 is related to the anthropogenic fire (e.g., López-Sánchez et al. 2017; Kosterov et al. 2020a) or the vegetation change (McClean et al. 2001). This is consistent with the environmental magnetic implication we discussed below.
If we take the IRM decomposition results into account, it suggests that the main magnetic components of the speleothem are magnetite. Comp-1 (Bh: 4.7–5.9 mT; DP: 0.35–0.38) can be identified as a low-coercivity multi-domain (MD) magnetite (e.g., Maxbauer et al. 2017; He and Pan 2020). Comp-2 (Bh:19.5–21.4 mT; DP: 0.31–0.37) represents a mixture of magnetite and maghemite (e.g., Abrajevitch and Kodama 2011; Font et al. 2014; Jaqueto et al. 2016). Comp-3 (Bh: 36.5–38.5 mT; DP: 0.24–0.29), which is observed in AAC-B3 and AAC-M, not in AAC-W3, are highly consistent with observations of oxidized pedogenic magnetite from previous work (e.g., Egli 2004; Maxbauer et al. 2017), which is considered as mainly single domain particles.
Thus, the assemblage of magnetic minerals in the grayish surface layer comprises a mixture of MD magnetite, PSD magnetite/maghemite and SD magnetite/maghemite. In contrast, the assemblage of magnetic minerals in the white inner part is dominated by MD magnetite and PSD magnetite/maghemite without pedogenic SD magnetite/maghemite. We suggest that the environmental condition that supplied magnetic minerals included in the surface grayish layer is different from that of the inner white part. Furthermore, hematite (or ε-Fe2O3) could be observed only in the reddish-brown layer, which was also recognized as a thin detectable feature in the SSM image of NRM after demagnetization and IRM acquisition.
FORC diagrams have three features: (1) the small peak around the Bc axis; (2) the spreading of the FORC distributions along the Bu axis with a sharp peak at around zero; (3) and a diagonal distribution extending to the lower-right between the Bu and Bc axis. These features can suggest the presence of pseudo-single domain (PSD) magnetite (Lascu et al. 2018). Alternatively, these features may suggest the mixture of single-domain (SD) and multi-domain (MD) particles, where SD particles are characterized as the existence of central-ridges around the Bu axis (Roberts et al. 2014) and multi-domain (MD) particles are represented as horizontal spreads (Lascu et al. 2018), respectively. Although FORC diagrams in this study have some noise even after stacking, we could tentatively interpret the FORC diagrams in combination with IRM component analyses that the gray part is dominantly composed of SD, PSD, and MD particles, whereas the white part is dominantly composed of PSD and MD particles. Recently, Ahmed and Maher (2018) showed that magnetite formation in Chinese paleosol occurs in well-drained, generally oxidizing soils. They also suggested that maghemite was present as an oxidized shell around the magnetite, where coatings of clay minerals are preventing the transport of oxygen to their surface. Their results have profound implications on our future studies in connecting to climate change and human activities.
Environmental implication
Results obtained from this study demonstrate that different magnetic properties can be detected clearly between the two layers using SSM. Concentrations of magnetic minerals are higher in the gray layer, whereas much smaller in the white part. The type of magnetic minerals contained in the two layers also can be distinguished, and a hiatus in growth appears to exist between the two layers.
According to Lascu and Feinberg (2011), the deposition rate of detrital magnetic minerals onto the speleothem surface is controlled by precipitation, flood frequency, and the surface soil above the cave. The depositional mechanism of gray layers and white part could potentially be attributed to changes in ambient environments. Previous studies showed that frequent events such as extreme climate changes, including floods and/or tsunamis, could have been recorded as distinct, several colored layers (e.g., Denniston and Luetscher 2017; Feinberg et al. 2020).
The distinct gray color on the white calcium carbonate speleothem may have been deposited with influence from non-carbonate minerals. The surface geology of Tongatapu island consists mainly of limestones overlain by two soil layers of andesitic volcanic ash, which are estimated to be aged at 20,000 and 5000 years from active volcanoes in the north, such as Tofua and Kao island (e.g., Cowie et al. 1991; Spennemann 1997; Manu et al. 2014). Soils of Ha' apai, Kingdom of Tonga, which are correlated with the soils of Tongatapu, are high in iron oxide minerals by weight (Childs and Wilson 1983). Furthermore, typical soils containing volcanic ash is acidic; thus, magnetite in the soil could oxidize to maghemite (Taylor and Schwertmann 1974). Therefore, the gray layers may have been formed due to volcanic eruptions 5000 years ago that occurred in the northern islands.
The other possible scenario would be the potential influence of human activities. The first human settlement on Tongatapu ranges from 2863 to 2835 cal BP (Burley et al. 2015). As the previous studies suggested, soil magnetic and archaeomagnetic studies can indicate that an increase in fire-related activity and an increase in the organic carbon supply from human activities, leading to a transformation into ferrimagnetic minerals (e.g.,Marmet et al. 1999; Hanesch et al. 2006; Fassbinder 2015). Thus, the results from this study also reveal similar magnetic properties that would point to the influence of anthropogenic fire use.
Moreover, AMS δ13C values of speleothem can provide supportive evidence for such human activity. Previous studies have suggested that variations in δ13C values of speleothems can be largely caused by two effects, both of which are potentially influenced by climate changes (e.g., Fairchild and Baker 2012a; Schwarcz 2013). When the climate is warm/arid, C4 plants with higher δ13C values are favored relative to C3 plants, leading to higher δ13C values in speleothems (e.g., Denniston et al. 2007; Zhang et al. 2015). On the other hand, sparse vegetation above caves leads to a lower proportion of organically derived and isotopically lighter CO2 dissolved in the drip water and results in contributing to higher speleothem δ13C values (e.g., Baldini et al. 2005; Genty et al. 2006). Carbon and nitrogen stable isotope analyses of bone collagen and apatite (Herrscher et al. 2018; Fenner et al. 2021) and pollen analysis in Tongatapu (Clark et al. 2015) indicated the development of horticulture ~ 2500 cal. BP for the supplement of the lack of marine foods. Therefore, the changes in δ13C values of the speleothem can suggest that the shift from dense to sparse vegetation was caused by primitive agriculture. In fact, Manu et al. (2014) reported that δ13C of repeatedly cropped soil is heavier than that of the primary forest soil with > 50% contribution of soil carbon from C4 grass vegetation in modern Tongatapu.
Future work, such as U/Th dating, is needed to discuss the processes affecting the magnetism of speleothems. The speleothem magnetism is still in the early developmental stages. However, by applying the method proposed in this study to other speleothems, speleothem magnetism can be a new paleoenvironmental proxy in addition to the conventional ones (e.g., oxygen isotope, trace elements/Ca ratio).