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Conspicuous granular and tubular cavities in volcanic glass are a widespread phenomenon in the oceanic crust, often attributed to microbial activity and etching (Furnes et al., 2008; Staudigel et al., 2008). These etch-marks are trace fossils, correctly referred to as ichnofossils (McLoughlin et al., 2007, 2009), after Greek ikhnos which translates to trace or track.

Already in 1986 Ross and Fisher (1986) reported etched grooves in glass shards from a Miocene tephra, which they compared to fungal borings in carbonate grains. During the 1990s, observations of granular and tubular microstructures were made in volcanic glass of drilled and dredged samples from the ocean floor (Thorseth et al., 1992, 1995; Furnes and Staudigel, 1999). Their biogenic origin were argued based on four criteria: (1) enrichments of C, N, and P in association with the etched cavities (2) C isotope values indicative of biologic activity (Furnes et al., 2001) (3) the presence of partially fossilized and encrusted microbial cells in the cavities, with forms and sizes matching those of the granular and tubular etch-marks, and (4) presence of nucleic acids and DNA from both bacteria and archaea in close association with the etch-marks (Giovannoni et al., 1996; Thorseth et al., 2001).

The direction of ichnofossils is perpendicular to a glass surface and inward into the glass. The granular type is made up of micron-sized, near-spherical cavities etched into the glass and later filled with authigenic minerals such as clays and Fe-oxides/hydroxides. The granular ichnofossils usually occur in clusters and are irregularly distributed on both sides of a fracture. The tubular ichnofossils consists of smoothly curved tubes, with varying morphologies including branching, spiraling, or segmented textures. Diameters are normally between one to a few microns, and ten to hundreds of microns in length. The traces usually occurs in assemblages irregularly distributed with respect to the opposing side of any cracks and are commonly associated with granular cavities. A comprehensive photographic atlas of the ichnofossils has been produced by Fisk and McLoughlin (2013).

The biogenicity of ichnofossils has lately been challenged, especially ichnofossils in ancient ophiolites and greenstone belts (Schopf, 2006; Grosch and McLoughlin, 2014). Alternative abiotic formation processes have been suggested involving the remains of fluid inclusion trails, radiation damage trails, or ambient inclusion trails (AITs) (Lepot et al., 2009; McLoughlin et al., 2010). One major issue is that the microorganisms responsible for the production of these structures have not yet been identified in modern micro-borings. Despite the on-going debate, establishing biogenicity of ichnofossils in volcanic glass is crucial for understanding early life on Earth but also since they may be potential biosignatures on Mars (McLoughlin et al., 2007, 2010; Sapers et al., 2014).

Encrusted or fossilized coccoidal microorganisms have been observed in dredged samples of seafloor-exposed basalt (Thorseth et al., 2003), and in drilled subseafloor basalt (Ivarsson et al., 2015c). Fossilized filamentous microorganisms have been identified in veins and vesicles in ocean floor basalts (Ivarsson et al., 2015c), and in ophiolites of Devonian age (Peckmann et al., 2008; Eickmann et al., 2009). They are both attributed to an endolithic lifestyle, using the walls of the open pore space for anchoring and subsequent outwards growth into the open voids. The fossils either occur as single microbial or fungal filaments or, more often, in large groups, filling the entire pore space of a vein system (Figure 1, inserted box, Figure 2, 3A–D). Their morphology can be highly varied, comprising curvi-linear segmented and branched filaments that may form complex mycelium-like networks. Sometimes they are characterized by central swellings, a turgid appearance or bean shaped structures. The diameters vary excessively, ranging between 2 and ∼30 μm with lengths between 10 μm and <1 mm. The thinnest filaments often have the most basic morphologies, resembling simple bacterial trichomes. The thicker filaments on the other hand, can vary in complexity, and usually have conspicuous septa-like segmentation (Ivarsson et al., 2015c).

The fossilized microorganisms comprise crypto-, chasmo-, and euendoliths and they can be divided into two main types, depending on preservation; (1) microorganisms enclosed and preserved in vein-filling minerals, commonly carbonate, or (2) fossilized communities mineralized and preserved directly into an open pore space (thus not entombed in secondary minerals). The first type consists of live microorganisms that have been instantly petrified in secondary mineral phases due to hydrothermal activity or alteration of the crust. They are usually permineralized by poorly crystalline clay phases and iron oxides with a high organic content, including elevated carbon levels, phosphates, hydrocarbons, lipids and occasionally even chitin (Ivarsson et al., 2012, 2015c). The second type, consist of microorganisms fossilized and mineralized in open systems where organic matter are easily oxidized and scavenged by other microorganisms. Thus, the second fossil-type rarely contains organics and are usually mineralized by a clay phase and/or iron oxides, similar to that of (1) above.

Microbial colonization occurs relatively early after cooling of the volcanic rock, presumably while fluid-flow is still active, and prior to substantial formation of secondary minerals. Colonization is initiated by the formation of a pioneer-type biofilm lining the pore space interiors (Figure 2A), which can be seen today as a mineralized crust of various hematite, hematite-carbon, and clay layers (Ivarsson et al., 2013, 2015a; Bengtson et al., 2014). The mineralized biofilm can be either smooth or botryoidal in appearance due to assimilated spherical and cell-like aggregates (Figure 3E,F). Filamentous structures protrude from the basal film into the open pore space, either solely, or more commonly in great abundance, and form complex, mycelium-like networks. Thus, the fossils are always rooted at the fracture- or vesicle-wall, where they became permineralized upon death (Figure 3G).

Filamentous fossils in the deep crust, have almost exclusively been interpreted as the remains of fungal hyphae, and when found in complex extensive networks, as fungal mycelia. They commonly have characteristic fungal morphologies such as repetitive septa, anastomoses between branches, a central pore and, though rarely, chitin in the cell walls (Ivarsson et al., 2012, 2015c). Associated conglomerates of coccoidal fossils with diameters of ∼10 μm have been recognized as yeast-like growth states (Figure 3E,F) and transitions between hyphae- and yeast-like stages are common. The fungal fossils also contain sporophores, fruiting bodies and resting structures like sclerotia. Based on morphological traits and the presence of chitin, fungi fossils from the Emperor Seamounts in the Pacific Ocean were interpreted as Ascomycetes or stem-group Dikarya (Ivarsson et al., 2012). Other fossilized fungi collected at the Vesteris Seamount, the Greenland Basin, were interpreted as Zygomycetes based on the presence of zygospores and various morphological stages of the Zygomycetes reproduction cycle (Ivarsson et al., 2015d). Molecular studies of extant fungi in subseafloor igneous crust are rare with the exception of one fungal isolate from drill cores from North Pond, Mid-Atlantic Ridge. The isolate was affiliated to the genus Exophiala of the order Chaetothyriales (Hirayama et al., 2015). Eight yeast-like fungal species belonging to Ascomycetes and Basidiomycetes were isolated from seafloor-exposed basalt at Vailulu’u Seamount, Pacific Ocean (Connell et al., 2009). Fungal strains of Ascomycetes and Basidiomycetes were also reported from vent fluids at the Lost City, the Mid-Atlantic Ridge (López-García et al., 2007), and have been isolated from marine sediments (Orsi et al., 2013a,b) including some of the deepest sediments on Earth at the Mariana Trench (Takami et al., 1997). Chytridiomycota has been detected in hydrothermal vents (Le Calvez et al., 2009) and methane seeps (Nagahama et al., 2011). Thus, the fungal diversity seems to be high in both marine sediments and in subseafloor-exposed basalts, with representatives from all major fungal divisions.

A close ecological association between heterotrophic fungi and chemoautotrophic prokaryotes has been identified in fossilized material. The symbiotic nature of this relationship may be a prerequisite for eukaryotic colonization and persistence in the subseafloor crust (Bengtson et al., 2014; Ivarsson et al., 2015c,a). Fossilized fungal mycelia have been shown to function as a framework on which syngenetic prokaryotic communities grow. Microstromatolitic Frutexites with a characteristic cauliflower-like appearance are often seen attached to hyphae, while cell-like structures forming suspended sheeths of strings in between hyphae may be affiliated with iron-oxidizing bacteria, such as Pyrodictium or Euryarchaeon SM1 (Figure 3C,D). Microstromatolites (Frutexites), have also been observed lining the walls of veins and vesicles, often in close association with fungal hyphae. Frutexites usually consist of laminated iron, or occasional manganese, oxides and silica-rich phases (Figure 4). The overall interpretation is that they are the result of iron- or manganese-oxidizing bacterial activity (Bengtson et al., 2014; Ivarsson et al., 2015c,a). Bacterial biofilms with Frutexites can be overgrown by fungal biofilms and mycelia that graze on the prokaryotic biomass, thus enabling fungal colonization of oligotrophic subseafloor voids (Ivarsson et al., 2015a).

The chemical and mineralogical composition of fossilized microorganisms in deep crustal habitats is often dominated by a poorly crystalline mix of Si, Fe, Al, Mg, and C with minor amounts of Na, K, Ti, and Ca, usually corresponding to smectite or montmorillonite-like clays (Ivarsson et al., 2015c), chamosite and illite (Peckmann et al., 2008). The clay permineralization is probably a result of microbial clay authigenesis (Peckmann et al., 2008). While filament exteriors are permineralized by clays, the interiors are usually characterized by iron oxides/oxyhydroxides (Bengtson et al., 2014; Ivarsson et al., 2015c,a). The carbon content of the fossils, ranging from ∼10 to 50 wt % C, varies greatly both between closely associated filaments as well as within the same filament. When the organic content is high, it may contain the presence of phosphates, hydrocarbons and lipids (Ivarsson et al., 2015c). Fossils have also successfully been stained by propidium iodide (PI), a dye that binds to cells with damaged cell membranes and traces of DNA (Ivarsson et al., 2008).

The fossilization process can be described as a transition in which the primary organic material of the organisms matures to a carbonaceous material prior to final mineralization by clays and Fe-oxides (Ivarsson et al., 2018). The negatively charged carbonaceous material attracts positively charged Si, Al, Mg, Fe (and minor Na, and Ca) cations dissolved in the fluids, which initiates subsequent adsorption and clay mineralization. Eventually, this can lead to complete mineralization by clays and Fe-oxides. Such a fossilization process is supported by the presence of partly mineralized fungal hyphae in deep granite settings (Drake et al., 2017). The fungal-precipitated clay may furthermore be altered to chlorite, if the fossils are subjected to metamorphosis (Bengtson et al., 2017).

The enrichment of C and the high degree of organic remains present in microbial fossils may be due to the distribution of elements within the fossils. Silicification is known to preserve microbial morphologies and cell content to some extent (Toporski et al., 2002). Metal cations, on the other hand, inhibit the autolysis (self-destruction of cells) and may thus preserve dead bacterial cells (Leduc et al., 1982). Ferric ions have been suggested to inhibit autolysis of bacterial cells and result in intact cellular preservation (Ferris et al., 1988). Iron is also known to form complexes, and stabilize organic compounds over geological time scales (Parenteau and Cady, 2010; Lalonde et al., 2012). Therefore, the high iron content of fossilized microorganisms in deep oceanic crust is most likely responsible for the high concentration of C and organic remains, such as lipids and chitin, found within many fossils.

The fossils in deep crystalline basement is not an exception and are, as all microfossils, tested against biogenicity criteria to exclude a possible abiotic explanation (Ivarsson, 2006). Filamentous biomorphs are known to form in laboratory experiments (Garcia-Ruiz et al., 2003), but still today, there are no reports of abiotic morphological equivalents produced within geological and geochemical conditions similar to that of subseafloor basalts. Microstromatolites occur on a morphological and structural continuum from bacterially mediated “shrubs” to abiotically precipitated dendrites (Chafetz and Guidry, 1999). The structures at the biological end of the spectrum are morphologically identical to Frutexites found in subseafloor environments (Bengtson et al., 2014), but it is important to stress that all fossil findings needs to be individually tested for biogenicity.

The most abundant rock type on Mars is mafic in composition with a volcanic origin, similar to terrestrial basalts (McSween and Treiman, 1998; Squyres et al., 2004; Poulet et al., 2009; Edwards et al., 2017). Most Martian meteorites have a distinct range in chemistry and are characterized by relatively Fe-rich, alkali-poor, undifferentiated basaltic rocks (Musselwhite et al., 2005; McSween et al., 2009). In situ investigations by Mars rovers have indicated a wider range in chemistry, showing a Fe-rich basaltic signature but with a more alkali-rich composition (McSween and Ruff, 2006; McSween et al., 2009). In Pre-Noachian and Noachian times, the martian basalts were subject to circulating groundwater. In places, they were covered by standing sea-like water-bodies, which permitted water percolation and fluid–rock interactions similar to ocean-covered basalts on Earth (Mustard et al., 2008; Ehlmann et al., 2009; Fassett and Head, 2011; Carter et al., 2013; Ehlmann and Edwards, 2014; Grotzinger et al., 2015; Goudge et al., 2016). Succeeding the Noachian eon, the “wet period” of Mars ended (Westall et al., 2015). However, parts of the martian bedrock were at times exposed to, and circulated by water due to ice melting associated with volcanic eruptions and meteorite impacts, occasional melting of ice at depths, and seasonal ice melting (Laskar et al., 2002; Tanaka et al., 2003; Velbel, 2012; Tornabene et al., 2013). Even though the martian surface was dry and cold the martian subsurface was subject to groundwater flow (Osinski et al., 2013; Grotzinger et al., 2015; Yen et al., 2017). Since the composition and alteration of martian basalts is similar to terrestrial counterparts (McSween and Ruff, 2006; Mustard et al., 2008; Ehlmann et al., 2009; McSween et al., 2009; Fassett and Head, 2011; Carter et al., 2013; Ehlmann and Edwards, 2014; Grotzinger et al., 2015; Goudge et al., 2016) it is reasonable to assume that any potential carbon-based life forms living in the deep martian rocks, would have the same geochemical and energetic prerequisites as life in basalts on Earth, and would be fossilized in a similar manner. If we assume the metabolic pathways and mode of life to be similar to that of life on Earth, the fossil record of volcanic rocks on Earth provide a valid analog to a potential martian fossil record.

The first step in the identification of possible in situ biosignatures, will be to optically recognize textures with a possible biological origin. Focus here should be on fossilized microbial colonies in open veins and vesicles, easily accessible for rover instrumentation. The NASA Mars mission 2020 will be equipped with 23 cameras, of which most cameras are for landing, driving and landscape monitoring, thus visualizing large geological features (Farley and Williford, 2017). Of specific interest for high-resolution investigations such as these, are the cameras MASTCAM-Z, SUPERCAM, PIXL, and WATSON on SHERLOC. MASTCAM-Z is a camera pair that provide color images and videos, three-dimensional stereo images, equipped with a powerful zoom lens. It has a pixel scale of ∼0.5 mm/pixel to ∼0.15 mm/pixel at 2 m, which is sufficient to resolve ∼1 mm features in the near-field (robotic arm workspace) and ∼3–4 cm features approximately 100 m away. SUPERCAM and PIXL are both combined instrument set-ups including cameras with resolution similar to MASTCAM-Z¹.

SHERLOC (The Scanning Habitable Environments with Raman and Luminescence for Organics and Chemicals) is an arm-mounted, Deep UV (DUV) resonance Raman and fluorescence spectrometer capable of analysis at <100 μm spot size. The laser is integrated to an autofocusing/scanning optical system, and co-boresighted to a context imager, which at the most typical use has a resolution of 15 μm/pixel corresponding to a spatial resolution of approximately 50–60 μm (Kenneth Williford, personal communication). In addition to the combined spectroscopic and macro imaging component, SHERLOC also includes a near-field-to-infinity imaging component called WATSON (Wide Angle Topographic Sensor for Operations and eNgineering), which enables science operations and imaging. WATSON captures the larger context images for the very detailed information that SHERLOC collects.

ExoMars Panoramic camera (PanCam) (Coates et al., 2012, 2017; Cousins et al., 2012; Yuen et al., 2013) will perform digital terrain mapping for the mission rover. It is a suite that consists of a fixed-focus, wide-angle, stereoscopic, color camera pair (WAC) complemented by a focusable, high-resolution, color camera (HRC). PanCam will enable characterization of the geological environment, from panoramic (tens of meters) to millimeter scale, at the sites the rover will visit. It will be used for detailed investigations of outcrops, rocks, and soils, and to image samples collected by the drill before they are delivered to the analytical laboratory. For more detailed investigations, the ExoMars is equipped with the close-up imager (CLUPI) that will obtain high-resolution color images for exploration of the depositional environments, but also detailed textural features in the form of morphological biosignatures such as biolamination (Josset et al., 2017). CLUPI will be mounted on the drill box and be equipped with several viewing modes, allowing the study of outcrops, the fines produced during drilling, and also imaging collected samples in high resolution before delivering them to the analytical laboratory. Imaging resolution varies with distance to target; it is, 8 μm/pixel at 11.5 cm distance with view area 2.0 cm × 1.4 cm, 39 μm/pixel at 50 cm distance with view area 10 cm × 7 cm, and 79 μm/pixel at 100 cm distance with view area 21 cm × 14 cm (Vago et al., 2017).

PanCam and CLUPI can be used simultaneously to visualize outcrops and rocks at progressively higher resolution. PanCam and CLUPI can also view drilling operations. As the drill penetrates into the ground, fines will be collected of which CLUPI can obtain high resolution images at 39 and 13 (drill lowered) μm/pixel (Josset et al., 2017). Furthermore, borehole wall data can be obtained by the Ma_MISS IR spectrometer (Mars multispectral imager for subsurface studies) (De Angelis et al., 2013; De Sanctis et al., 2017), integrated in the drill tip. The PanCam HRC and CLUPI are capable of imaging samples placed by the drill on the Core Sample Transport Mechanism (CSTM), and finally CLUPI can obtain high-resolution images (∼8 μm/pixel) of the sample.

Resolution of the payload cameras controls the possible detection limit. Therefore, because of size limitations, identification and characterization of individual filaments and cell structures will not be possible with the current optical instrumentation. This also includes the detection of narrow ichnofossils and similar microbial traces. Instead, an approach involving structures <50–60 μm have to be developed. This includes fossilized communities with an internal morphological variation that may be detectable, or with a diagnostic mineralogy/geochemistry. Examples of possible targets are three-dimensional fungal mycelium with diameters exceeding 1 mm (Figure 5A). These mycelium often form a network of filamentous and/or yeast cells that grow perpendicular and markedly from a mineral or host rock surface. Another target can for example be extensive microstromatolites with well-defined and visible laminations that vary in color due to differences in mineralogy.

Figure 5A shows a vesicle ∼5 mm in diameter, with fossilized fungal mycelium lining the walls. Microstromatolitic Frutexites, ∼30–100 μm in diameter, have grown and become fossilized on the fungal hyphae as orange or yellow accumulations. Figure 5B is modified to 8 μm/pixel, the resolution of CLUPI on ExoMars, indicating a decrease in resolution compared with the original image. It is, however, still high enough resolution for details like hyphae and even Frutexites to be identified. Figure 5C is modified to 15 μm/pixel, the best possible resolution of SHERLOCK on NASAs Mars mission 2020 (Kenneth Williford, personal communication), and shows a substantially decreased resolution and image quality, which makes identification of individual hyphae or microstromatolites difficult. It is, however, possible to see that the vesicle in whole carries interesting morphological features along the vein walls that should be analyzed in situ for mineralogy and organic content or targeted for later Mars sample return.

During mineralization and fossilization of microorganisms, organic remains are incorporated and bound by mineral phases, especially iron oxides and clays. If the fossils only experience moderate taphonomic alterations during diagenesis, the organic content will remain more or less unaltered and easy to detect by in situ measurements. The organic content can be used as a strong argument when proving biogenicity of a putative microfossil or biomineral, but also, if complex biomolecules are preserved as biomarkers, to constrain the taxonomic affiliation of a fossil organism. Lipids and proteins, for instance, can be preserved during long periods of time and are usually specific for organismal affinity.

As discussed above, the optical system on SHERLOC is coupled with a Deep UV (DUV) resonance Raman and fluorescence spectrometer, utilizing a 248.6-nm DUV laser with <100 μm spot size. Deep UV-induced native fluorescence is very sensitive to condensed carbon and aromatic organics, with a detection-capacity at, or below, 10–6 w/w (1 ppm) at <100 μm. SHERLOC’s deep UV resonance Raman therefore enables detection and classification of aromatic and aliphatic organics with sensitivities of 10–2 to below 10–4 w/w at <100 μm spatial scales. Additionally, the deep UV Raman allows detection and classification of aqueous minerals with grain sizes below 20 μm.

The fungal hyphae and the Frutexites in Figure 5 have been analyzed by Raman and shown to contain carbonaceous matter (Bengtson et al., 2014). The spot size for these analysis were ∼1–3 μm, thus the analysis were done with higher precision than the instrumentation coupled to SHERLOCK will be. However, with a <100 μm spot size the overall mineralogy will be analyzed and give a composite spectrum including clays, iron oxides and carbonaceous matter, thus a comparable sample on Mars would be possible to analyze and detect organic remains within.

Pancam on ExoMars is coaligned with an IR-spectrometer (ISEM) that enables identification of mineral phases in bulk rock- or soil samples at a distance. As mentioned above, the Ma_MISS is a miniaturized IR spectrometer integrated in the drill tool for imaging of the borehole wall as the drill is being operated. Ma_MISS has a spectral range of 0.4–2.2 mm with spatial resolution of 120 μm. CLUPI, however, is not coupled with any instruments for chemical detection. When samples have been selected by Pancam, ISEM, and CLUPI, they are deposited in the core sample transport mechanism (CSTM) where PanCam HRC and CLUPI can image the sample during a few minutes. After the imaging exercise, the CSTM move the sample into the analytical laboratory where a rock crusher produce particulate matter having a Gaussian size distribution, with a median value of 250 μm. The implications of this maneuver is that any identified morphological structures of interest will be destroyed, and correlation between morphological biosignatures and molecular fossils will thus thereafter be impossible.

Microorganisms continually interact with their environment through their metabolism, which itself fosters mineralization and may favor fossilization. Secondary minerals resulting from metabolic activity are commonly of intermediate to high specificity for the actual metabolism. Detailed in situ analysis at a micro scale can be performed on mineralized fossils and associated biominerals by laser ablation inductively coupled mass spectrometry (LA-ICP-MS). Such instruments are optimized to analyze isotopic contents of individual microorganisms and microbial associations on a micro-scale. Examining the elemental and isotopic composition of the mineralized microbial remains, will allow a constraint of primary and secondary metabolites and bioavailable energy sources, information that can be used to determine the biological affinity of microfossils. For example, δ¹³C values of authigenic calcite will indicate the presence, or past presence of anaerobic methane oxidation, or methanogenesis, while δ³⁴S values in pyrite will indicate potential fractionation due to bacterial sulfate reduction (Drake et al., 2015, 2017). Isotopic investigations will therefore help to constrain metabolic pathways and can clarify the microbial involvement in the cycling of elements between rocks, microorganisms and water.

Miniaturized laser ablation ionization mass spectrometry (LIMS)-time-of-flight mass spectrometer for space missions capable of performing in situ isotope measurements of minerals at planet surfaces in vacuum have been developed and tested, but not included, in the payloads of ExoMars or NASAs mars mission 2020 (Wiesendanger et al., 2018). Isotope measurement thus have to be targeted for the subsequent Mars sample return mission.

Drill cores and other kind of samples retrieved from Mars offer the opportunity of a considerably more detailed chemical and morphological sample investigation. As long as the payload instrumentations enable sampling of material with a high potential for biosignature preservation, the chances of identifying and characterizing morphological and/or molecular fossils in a piece of rock is therefore higher by magnitudes in laboratories on Earth compared to in situ on the surface of Mars.

Microimaging techniques are central in all microfossil investigations, in order to resolve microbial morphology, variations in composition and relation to host rock and associated mineral phases. Methods and instrumentation can vary depending on access to equipment and facilities. For example, synchrotron-based X-ray tomographic microscopy (SRXTM) is a non-invasive technique that allows the study of rock and fossil structures in 3 dimensions, which has been decisive in the identification of deep-biosphere fossils of different ages (Ivarsson et al., 2012, 2013; Bengtson et al., 2014, Bengtson et al., 2015c,a). Recent identification of fungus-like fossils from 2.4 Ga vesicular basalts obtained by SRXTM (Bengtson et al., 2017) were instrumental in distinguishing pore-dwelling cryptoendoliths from rock-boring euendoliths. The physical interaction between microorganisms and minerals represented by either biomineralization or bioerosion structures have thus been successfully visualized and studied by SRXTM (Bengtson et al., 2014; Ivarsson et al., 2015c,a). So far, most tomographic work on deep biosphere fossils have been executed at the micro-tomography beamline TOMCAT at the Swiss Light Source synchrotron facility, Paul Scherrer Institute, Switzerland. The TOMCAT beamline has provided tomographic datasets with micron-sized sample resolution, which has resulted in increased insights into the spatial correlation between microbes and vein filling minerals as well as their association with vesicle/fracture surfaces within the basalt. In addition, new facilities such as the synchrotron Max IV in Lund, in particular the beamlines DanMAX and NanoMAX, will have the capacity to resolve nanometer-scale features, which will open a new path toward the investigation of cryptic microfossils and their ultrastructural features.

For the examination of volcanic biosignatures and fossils, initial mineralogical and fossil identification should be performed using a combination of light and fluorescence microscopy, scanning electron microscopy (SEM) coupled with energy dispersive X-ray spectroscopy (EDS), X-ray diffraction (XRD), and X-ray fluorescence (XRF). Further detection and characterization of organic material in fossils may be targeted by Raman spectroscopy and Time-of-Flight secondary ion mass spectrometry (ToF-SIMS) (Ivarsson et al., 2015c).

The above analysis should also be complemented by bulk analyses. The extraction/characterization of biomarkers may be performed using gas chromatography (GC), coupled gas chromatography/mass spectrometry (GC/MS), and coupled high-pressure liquid chromatography/mass spectrometry (HPLC/MS). For compound-specific stable carbon isotope compositions, gas chromatography-isotope ratio monitoring-mass spectrometry (GC-irm-MS) can be used.

To establish an atlas of volcanic-based microfossils for the upcoming Mars missions, we need to enhance the current knowledge of fossilized microorganisms in deep igneous rock and in deep time, as reviewed above. The ambition should be to establish a fossil record from the present and backward, to trace microbial components of the deep biosphere in time, and to establish an improved framework for the timing of major evolutionary lineages. The main goal is to apply this knowledge to future exploration missions on Mars, with the possibility of identifying extra-terrestrial fossilized microorganisms from volcanic habitats.