Introduction

Frutexites are microstromatolitic structures with a purported biological origin, and have been reported from a variety of subsurface environments such as deep-sea basalts (Bengtson et al. 2014), granitic rocks (Heim et al. 2017), and marine carbonate caves (Guido et al. 2016). Frutexites display fine-scale lamination and a directional growth orientation forming columnar structures and are characteristically composed of iron oxides and/or manganese oxides and carbonates (Maslov 1960; Horodyski 1975; Myrow and Coniglio 1991, Heim et al. 2017). As noted by Jakubowicz et al. (2014) and Guido et al. (2016), Frutexites are typically restricted to environments with low light availability and low sedimentation, which is consistent within cryptic deep subsurface environments.

Although Frutexites are generally implicated as products linked to direct or indirect biological processes, few reports exist of their immediate connection to microbial activity. Of these reports, Horodyski (1975) and Chafetz et al. (1998) reported microfossils in proximal association with Frutexites-like structures and interpreted their formation as a product of biologically-induced precipitation. Recent work by Heim et al. (2017) further explored the microbial connection and identified active biofilm communities growing in association with modern Frutexites-like structures, which were shown to contain nitrogen- and iron oxidizing bacteria as well as Archaea and fungal biomarkers. Other studies have shown the close connection between organic matter and trace metals, where dissolved organic matter in aqueous environments forms complexes with trace metal ions (Hirose 2007; Rodriguez-Martinez et al. 2011). Organic matter such as cells and extracellular polymeric substances have a negatively charged surface that attract positively charged metal cations (i.e., Mn2+ and Mn3+) to form ligands and complexes (Frankel and Bazylinski 2003; Sjöberg et al. 2020). Adsorbed metal cations may eventually become hydrolyzed and oxidized resulting in the precipitation of metal oxides (Frankel and Bazylinski 2003), which is a process that has been implicated in the formation of Frutexite-like structures (Guido et al. 2016).

In this study, we have analyzed Frutexite-like structures composed of Mn-oxides occurring in deep subsurface ultramafic rock hosted fractures from the Mid-Atlantic ridge. Identifying fossilized microorganisms and associated structures, such as Frutexites, in the geologic record involves detailed analyses of micrometer-scale features and the ability to differentiate between biological and abiotic origins of often cryptic signatures (Benzerara and Menguy 2009). Multiple lines of evidence are typically required to determine whether a putative fossil is biogenic, including identifying characteristic morphological and chemical signatures, and establishing the overall habitability of the environment in which the fossil is found (McLoughlin et al. 2007). To explore the potential biogenicity of the Frutexites-like and microfossil structures reported here, we performed in situ chemical and isotopic analysis and found evidence consistent with a biological origin.

Samples and methods

The 1272A 2R1 40–45 cm sample was collected from the slow-spreading Mid-Atlantic Ridge (MAR) during Ocean Drilling Program (ODP) Leg 209 at drill hole site 1272A (Alt et al. 2007; Bach et al. 2004). The The sample was obtained from a depth of 13.3 m below seafloor (mbsf) at a seafloor depth of 2560 m below sea level. The host rocks consist of serpentinized harzburgite with an approximate age of < 1 Myrs (Alt et al. 2007; Bach et al. 2004). The sample are characterized by aragonite veins cutting through the serpentinized host rock containing putative morphological biosignatures in the form of microstromatolites embedded in the aragonite matrix (Ivarsson et al. 2018). Doubly polished thin sections were made from all samples to identify fossilized microorganisms and minerals by optical microscopy.

Scanning electron microscopy

Element distribution and imaging were performed using two different ESEM´s. The first ESEM is a XL30 Scanning Electron Microscope with a field emission gun (XL30 ESEM-FEG). The ESEM was equipped with an Oxford x-act energy dispersive spectrometer (EDS), backscatter electron detector (BSE) and a secondary electron detector (SE). The second instrument is a Zeiss Supra 35VP (Carl Zeiss SMT, Oberkochen, Germany) field emission SEM, equipped with a VPSE detector for low vacuum conditions, and a Robison BSD for back scatter detection. Also equipped with a EDAX Apex 4 (Ametekh, Mahwah, USA) EDS-detector for X-ray microanalysis. None of the samples were coated.

Raman spectroscopy

Identification of mineral phases and carbon phases were made on the thin sections using a Raman spectrometer using a multichannel Dilor XY spectrometer with an incident laser beam of 514.5 nm and an Innova 70 Ar laser. A Si standard wafer (111) with a Raman peak at 520.4 cm−1 were measured daily as a reference prior to the sample analyses. Peak and element analyses were made using INCA Suite 4.11 software. The laser power was determined experimentally but did not exceed 0.5 mW on the organic carbon analyses and the manganese oxides (Fig. 2).

Ion microprobe analyses

In situ carbon isotope analyses were made using the NORDSIM Ion Microprobe (Cameca IMS 1280) at the Department of Geosciences, Swedish Museum of Natural History. The polished samples were mounted on a stainless-steel sample holder and examined with SEM prior to the Microprobe analysis. The SIMS-analysis (10 μm lateral beam dimension, 1–2 μm depth dimension) of carbon isotopes in carbonate and the microstromatolites was performed on a CAMECA IMS1280 ion microprobe. Data were normalized for instrumental mass fractionation using matrix-matched reference materials (S0161, see Supplementary for more details) which were mounted together with the sample mounts and analysed after every sixth sample analysis. Results are reported as per mil (‰) δ13C based on the PeeDee Belemnite (PDB)-reference value. In total, 18 analyses were made for δ13C (12C/13C) of the sample, divided into two analytical sessions (Fig. 3a and b). The reference material was measured to + 0.22 ± 0.40‰.

Results

Mineralogy

Sample 1272A 2R1 40–45 cm is composed of serpentine cross-cut by aragonite veins (Alt et al. 2007 and Fig. 1). Spot analyses using SEM showed that the dominating elemental composition was O, Mg, Si and minor Ca and C, whereas the vein composition was dominated by O, Ca and C. Raman spectra of the vein mineralogy is dominated by a major peak at 1085 cm−1 and smaller peaks at 155, 207, 706 cm−1 corresponding to the reference RRUFF ID R080142 of aragonite (Fig. 1). A few of the matrix analyses showed sharp spectral peaks at 1082 cm–1 and minor peaks at 153, 180, 209 and 278 cm−1 corresponding to the reference RRUFF ID R050128 of calcite. However, the dominating type of carbonate is aragonite, which has been reported previously (Ivarsson et al. 2018). Black, branching structures in the vein walls was dominated by Raman spectra with only one single peak at 675 cm−1, which can be assigned to the strongest band in magnetite (Fe2O4) for green laser (Otero-Lorenzo et al. 2014).

Fig. 1
figure 1

Overview of the thin section sample (a) and the overall composition of the matrix where b shows the composition of the aragonite vein, the branched magnetite structures with the included Raman spectra and c the EDS spot analyses of the bulk serpentine and aragonite veins

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Microstromatolites

At the interface between the host rock and the vein cauliflower-like microscopic structures (Fig. 2a), a few hundred µm in length and up to 50 µm in diameter protrude from the vein walls in a perpendicular fashion and are embedded in the aragonite (Fig. 2a). These structures branch frequently and have an opaque appearance under reflected light microscopy (Fig. 2b). Optical microscopy with a Nikon P-CL 530 nm retardation plate show the structures to consist of dark, magenta colored anisotropic minerals (Nobile et al. 2009) in contrast to the yellowish crystalline aragonite and serpentine (Fig. 2c). EDS analyses (Fig. 2d) and elemental mapping of the sample shows a clear difference between the elemental composition of the microstromatolites (MS) and the rest of the sample (Fig. 2e–h). The vein walls are composed of Mg, Si, O and Fe whereas the MS is dominated by Mn and O. In cross section, the structures exhibit directional growth and apparent successive layers (Fig. 21–l).

Fig. 2
figure 2

Overview over the MS´s where a shows optical microscopy image of the MS growing perpendicular to the vein walls (marked by dashed lines), b shows the MS under reflected light, c shows the anisotrophy of the MS using P-CL retardation plate, d shows the EDS image of the MS, e–h the EDS mapping of the MS and vein walls and i–l shows the layered structures of the MS and where the arrows in k marks several of those layers

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Raman spectra acquired from the MS (Fig. 3a) showed prominent spectral peaks where the D1 band at ~ 1350 cm−1 and the D2 band at 2920–2962 cm–1 is visible together with the G-band at ~ 1580 cm–1 (Fig. 3b) corresponding to organic carbon (RRUFF ID R120025) and often displaying some degree of fluorescence. Several of the prominent Raman peaks in the MS spectra may be attributed to various Mn phases and species, such as the relatively large and sharp peak at 636 cm−1, which can be ascribed to todorokite (RRUFF ID R0707257, Fig. 3c) and anhydrous manganese oxide [((Na, Ca, K, Ba, Sr)1−x(Mn, Mg, Al)6O12·3-4H2O)]), Fig. 3d. The MS was dominated by manganese oxides, which was evident EDS analyses (Fig. 3c, d) and in the Raman spectra that showed peaks at 580 cm−1, which can be assigned to octrahedrally coordinated Mn3+ (Bernardini et al. 2021) and/or todorokite (Ostrooumov 2017). A few peaks coincided with sharp spectral peaks at 1082 cm–1 and minor peaks at 153, 180, 209 and 278 cm−1 corresponding to the reference RRUFF ID R050128 of calcite.

Fig. 3
figure 3

Raman spectra of the MS where a show Bernardini et al. (2021) reflected light image of the MS, b shows manganese spectral bands coinciding with carbonaceous matter b but is dominated by various types of manganese oxide spectra such as todorokite (c) and octrahedrally coordinated Mn3+ (d)

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SIMS analyses of putative fossil microstromatolites

In-situ δ13C analyses on the CM using Ion Microprobe (NORDSIMS) yielded an average value of – 35.4 ± 0.50 ‰ (n = 7, Fig. 4, Supplementary Table 1) compared with the crystalline calcite standard S0161 (Fig. 4), while analyses of the aragonite matrix yielded an average δ13C value of + 0.67 ± 0.40 ‰ (n = 13, Fig. 4), resulting in a Δ13C of 36.07 ± 0.64 ‰. Though the absolute δ13C values reported here likely reflect differences in ionization efficiencies owing to matrix effects, the offset in carbon isotopic composition between the MS structures and the aragonite matrix are clearly significant and correspond to an average offset of − 36.07 ± 0.64 ‰.

Fig. 4
figure 4

SEM backscatter image a showing the ion microprobe beam spots and the corresponding δ13C values of the aragonite matrix and b the ion microprobe image with corresponding δ13C ion microprobe measurements of the MS and matrix. The dashed line marks the MS structures. MS microstromatolite structure, ARA aragonite, SRP serpentinite

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Discussion

The fossil record of igneous crust

The deep biosphere (DB) includes organisms living within sediments and igneous rock in both oceanic and terrestrial crust. The DB is estimated to be the second largest reservoir of biomass on Earth representing at least one-tenth to one-third of Earth’s total biomass, and is understood to mostly consist of prokaryotic microbial life but also eukaryotes, i.e., fungi (Bengtson et al. 2014; Magnabosco et al. 2018). The deep biosphere is still poorly understood, but recent advances in genomic sequencing coupled with the availability of new samples have revealed a surprising diversity of this hidden realm. It is clear that the DB is a substantial reservoir of novel microbial phyla (Magnabosco et al. 2018) and that the interplay between the DB and the surface environments is an important process to all life on Earth. By studying the deep subsurface fossil record, we may further investigate the evolution and distribution of early life on Earth and add to our understanding the interplay between the geosphere and biosphere through time. There is a growing literature of paleo-rock-hosted life from oceanic to terrestrial deep environments and an increased understanding of the biochemical energy flows within this biosphere (Ivarsson et al. 2020).

Microorganisms including fungi are known to colonize rock surfaces, inhabit pore spaces, and utilize the energy released either through direct microbial alteration of rock surfaces or by using the products of abiotic water/rock interactions. Fossilization of these microbes through sudden precipitation of minerals, such as carbonate or microcrystalline quartz, entombs the microbial colonies in situ. Thus, the fossil rock record not only provides information on the abundance and type of life that inhabits the DB, but also to the extent of their habitable niche environments including fractures, pore spaces and veins within the rock matrix. Raman analyses of the vein filling material surrounding the MS showed that the majority of the minerals was composed of aragonite (Fig. 2). The MS was dominated by manganese oxides, which was evident both in EDS analyses (Fig. 1e–h) and in the Raman spectra (Fig. 3) that showed peaks at 580 cm−1, which can be assigned to octrahedrally coordinated Mn3+ (Bernardini et al. 2021) and/or todorokite (Ostrooumov 2017), suggesting an oxidizing precipitation environment. Several of the prominent Raman peaks in the MS spectra may be attributed to various Mn phases and species, such as the relatively large and sharp peak at 636 cm−1, which can be ascribed to todorokite and anhydrous manganese oxide. The Mn is present only in the MS and absent in the surrounding rocks (Fig. 2e–h), indicating that the MS have formed through accumulation of Mn from the fracture fluids. Manganese is common in mafic and ultramafic rocks, constituting approximately 0.16% of the total elemental abundance (Cannon et al. 2017) and have likely been leached to the serpentine fluids during the hydration of the rocks, and subsequently accumulated in the MS. No manganese oxide peaks were found with the Raman analyses other than within the MS. The MS often display broad Raman peaks and sometimes fluorescence, consistent with an amorphous mineralogical composition as well as the presence of organic material and metal oxides (Olcott and Marshall 2015). This anisotropic nature of the MS´s are confirmed by analyses using optical microscopy retardation plate (530 nm, Nobile et al. 2009) showing strong anisotropic colours of the MS, in contrast to the surrounding crystalline aragonite (Fig. 2c). Even though anisotropy is not a biosignature, it is indicative of a different precipitation environment than the closely precipitated branched magnetite or the surrounding crystalline aragonite (Fig. 1). The Raman analyses of the MS showed multiple peaks coinciding with carbonaceous matter (CM), which may have been derived from decomposed biofilms, but that is not possible to discern using 514 nm Raman spectroscopy. The D and G bands are usually used to detect CM of geological samples of various origins, such as microbialites, Frutexites and older fossils (Guido et al. 2018, 2019; Allwood et al. 2006; Greco et al. 2018). Close to the vein walls, black mineral structures roughly resembling the MS, were all composed of pure magnetite, likely a product of serpentinization (Fig. 2). Our samples are found in an aragonite vein in an ultramafic oceanic rock that has experienced partial serpentinization (i.e., the hydration of Fe/Mg-bearing minerals and formation of serpentine phyllosilicate phases). Serpentinization is an exothermic reaction capable of releasing a wide range of energy sources for microbial growth, such as H2, Fe2+/Fe3+, and other metals, and the deep biosphere is to a large extent sustained by the products from serpentinization (Klein et al. 2015; Schrenk et al. 2013).

Biogenicity

To test the biogenicity of the MS, the following criteria was used: geological context and associated habitability, morphology, and chemical (i.e., elemental and isotopic) composition. Filamentous fossils have been described in the same sample in a previous study (Ivarsson et al. 2018) where the biogenicity criteria was established for the filaments within the same geological context. The mineral succession with enclosing aragonite indicates that the formation of the MS preceded the precipitation of aragonite and thus grew on the fracture walls while the vein still was open and fluids or seawater circulated the system. This is supported by the lack of alteration rims of the contact between the aragonite and MS. Moreover, the MS do not seem to follow any crystal planes or microcavities, indicating that the MS preceded the aragonite. A dissolution of pre-existing aragonite would have produced an alteration rim surrounding the MS, which was not visible in our samples. As the aragonite formed and sealed the veins, the MS were embedded and preserved in situ. According to Alt et al. (2007), the δ18O signature of rocks from a nearby drill hole 1272A 2R1 40 are consistent with low-T serpentinization at under 150 °C. Bach et al. (2004) described the presence of the mineral iowaite, which also suggests a low-T regime, forming from Fe-rich brucites during later stages of serpentinization (Bach et al. 2004). Reported formation T < 14 °C (determined using oxygen isotopes) for the aragonite enclosing the MS shows that the T at the time of growth was well within the T range for biological activity (Bach et al. 2004).

The morphology of the microstromatolites is characterized by apparent alternating laminations (Fig. 1i–l), fractal branching growth (Fig. 1), and a distinct direction of growth, which is consistent with previously reported microstromatolites, which have been interpreted as having a biological origin (Guido et al. 2016; Heim et al. 2017). The size distribution of MS reported in these environments are similar to the structures we describe here (Heim et al. 2017; Böhm and Brachert 1993; Myrow and Coniglio 1991). Similar morphologies to the MS´s have been reported by Sjöberg (2019, Fig. 5) in the Ytterby mine in Sweden, where active microbial communities create similar morphologies to the MS and where induced concentration of elements precipitates birnessite-type Mn oxides. Bio-mediated apparent layering, wad like and reticulated morphologies resembling those in our study is also reported in these open mine systems (Fig. 5). Even though abiotic oxidation of Mn2+ is possible in such aerobic and slightly basic (pH 8) environments, it is a considerably slower process than a microbially mediated oxidation of Mn2+ to Mn3+ (Sjöberg 2019). In mafic/ultramafic systems, fluid pH values are commonly above 8, and at such conditions dissolved Mn2+ is partly hydrolyzed, easily adsorbed, and locally concentrated to any available solid surfaces, such as vein walls, cells etc. The trapped Mn2+ can then be further oxidized by epilithic or endolithic microorganisms, forming the characteristic morphologies seen in Fig. 5.

Fig. 5
figure 5

EDS images showing the morphological similarities between recent, bio-mediated manganese oxides (inset) with its putative laminations (black arrows) and the inner, wad like and reticulated grooves (red arrows). The inset image is modified with permission from Sjöberg 2019 showing manganese oxides mediated by biology in recent natural precipitates on mine walls in Ytterby mine, Sweden. (Color figure online)

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These fractally branching bodies are morphologically also very similar to structures that have been previously reported in the literature and identified as Frutexites, which are found in cryptic habitats and hydrothermal vents through the Proterozoic and Phanerozoic (Walter and Awramik 1979; Böhm and Brachert 1993) and within subseafloor igneous environments (Bengtson et al. 2014). The nature of microstromatolites has been the subject of continual discussion, with most interpretations in the literature indicating a bacterial origin, however abiotic chemical precipitation has also been considered as an alternative. This question reflects the well-known problem of how to distinguish biogenic stromatolites from non-biogenic precipitates (Grotzinger and Rothman 1996). Microstromatolitic precipitates in hydrothermal environments are found to range in morphology and structure from bacterially-mediated ‘shrubs’ to abiotically precipitated dendrites (Chafetz and Guidry 1999). The microstromatolites in sample 1272A 2R1 40–45 are consistent with the first, bacterially formed, category, while microstromatolite structures in the literature ascribed as abiotic are largely dominated by euhedral crystalline forms (Braissant et al. 2003).

Because morphology is merely an indication of biogenicity, additional chemical information of the structures is required to establish a biologic origin. The CM detected by Raman spectroscopy is consistent with biologically derived C in the MS since both the D1 and D2 bands were present (Fig. 3). However, because the presence of D and G Raman bands are only evidence of disordered C and may not represent a definitive biosignature (Allwood et al. 2006), bulk and in-situ δ13C was also measured. Our in-situ δ13C measurements of the MS yielded an average value of – 35.4 ± 0.50 ‰, while the surrounding aragonite matrix yielded an average value of δ13C = 0.67 ± 0.40 ‰, consistent with a marine carbonate (Emerson and Hedges 2008). Assuming fractionation proceeded from a common reservoir of dissolved inorganic carbon, the offset between average δ13C values of CM and aragonite is − 36.07 ± 0.90 ‰, which is consistent with fractionation induced by metabolic processes such as the Wood-Ljungdahl pathway or dissimilatory sulphate reduction (Detmers et al. 2001; Londry and Marais 2003). Sulphur isotopes (δ34S) previously measured in the samples had values of approximately – 32‰, which were interpreted as resulting from microbial reduction of seawater sulphate via sulphate reducing bacteria (SRB, Alt et al. 2007). Abiotic sulphate reduction was excluded due to kinetic and mass balance mismatches with abiotic processes (Alt et al. 2007). Characteristic fractionation patterns of δ13C are often cited as strong evidence for biological activity (Schopf, 1993; Ueno et al. 2004), however this interpretation has been challenged by some authors (McCollom and Seewald, 2006). While δ13C alone may not be strictly sufficient for claiming a biologic origin for microbial structures, a combination of several biologic signatures would provide stronger evidence. SRB have been suggested together with other bacteria such as Fe-reducing bacteria to form microstromatolites as a result of biofilm mineralization under fluctuating environmental conditions (Krumbein 1983). Based on elemental composition and observed δ13C and δ34S isotopic fractionation, we suggest a biological origin of the CM matter within the microstromatolites, and furthermore suggest that signatures are consistent with SRB activity. Energy sources available for microbial colonization and growth are likely derived from serpentinisation reactions where oxidation of reduced Fe species (such as pyroxene and olivine) may reduce water to H2 through mineral–water interactions. Hydrothermal venting and active serpentinisation is common in the sampled area, with elevated concentrations of both CH4 and H2 having been reported (Alt et al. 2007). In combination, the morphological features, presence of carbonaceous material, geologic context favourable for life, in-situ measurements of δ34S indicative of SRB activity, and in situ δ13C measured on the MS thus indicates a biologic origin of our microstromatolites (Benzerara and Menguy 2009).

Conclusions

We analysed samples from the Mid-Atlantic-Ridge collected during IODP Leg 209, and found microscopic structures that resembled fossilized Frutexites. Observations conducted via microscopic imaging, Raman spectroscopy, and in-situ SIMS carbon isotope analyses indicate a biological origin for these structures. Frutexites are typically reported as primarily composed of iron oxides, thus our results broaden our understanding of Frutexites to include biogenic structures ofcomposed dominantly of Mn oxides. Additionally, the structures in our samples are found in an ultramafic environment, which is not commonly correlated with Frutexites. Combined, our data suggest that the microbial colonization of the deep subsurface is widespread and that similar structures and morphologies can be found in a vast number of different environments.