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1 Introduction

Mars orbits close enough to Earth to have benefitted from a rich history of robotic exploration and discovery. It is thus the sole example we can justifiably invoke to illustrate the sustained effort required to search for signs of life elsewhere.

Based on what we knew about planetary evolution in the 1970s, many scientists regarded as plausible the presence of simple microorganisms on other planets, especially if their surface conditions could sustain at least episodically the presence of liquid water—which had been found not to be the case on Venus or the Moon. The 1976 Viking landers can be considered the first missions with a serious chance of discovering signs of life on Mars. That the landers did not provide conclusive evidence was not due to a lack of careful preparation. The Viking results were a consequence of the manner in which the question of life detection was posed, seeking to elicit signs of microbial activity from potential extant ecosystems within the samples analyzed (Klein et al. 1976). The twin Viking landers conducted the first in situ measurements on the martian surface. Their biology package contained three experiments, all looking for indications of metabolism in soil samples (Klein et al. 1976). One of them, the Labeled-Release Experiment, produced very provocative results (Levin and Straat 2016). If other information had not also been obtained, these data would have been interpreted as proof of biological activity. However, theoretical modelling of the martian atmosphere and regolith chemistry hinted at the existence of powerful oxidants that could, more or less, account for the results of the three biology package experiments (Klein 1999). The biggest blow was the failure of the gas chromatograph mass spectrometer (GCMS) to acquire evidence of organic molecules at the parts-per-billion level. With few exceptions, the majority of the scientific community concluded that the Viking findings did not demonstrate the presence of extant life (Klein 1998, 1999). As a consequence, our neighbour planet lost much of its allure and a multi-year gap in Mars exploration ensued.

During the 1990s and early 2000s, orbiters (Malin and Edgett 2000), landers, and the very successful Mars Exploration Rovers (Squyres et al. 2004a, b) focused on surface geology—searching for signs of life was not part of their objectives. However, this would change after findings by Mars Express 2003 and Mars Reconnaissance Orbiter 2005 revealed many instances of finely layered deposits containing phyllosilicate minerals that could only have formed in the presence of liquid water, thus reinforcing the hypothesis that early Mars had been wetter than today (Poulet et al. 2005; Bibring et al. 2006; Loizeau et al. 2010, 2012; Ehlmann et al. 2011; Bishop et al. 2013, 2018; Michalski et al. 2013). These discoveries rekindled the interest in young Mars as a potential abode for life.

More recently, two missions that have improved our understanding of chemical conditions on the martian surface are the 2007 Phoenix lander and the 2011 Curiosity rover. Phoenix included, for the first time, a wet chemistry analysis instrument that detected the presence of the perchlorate (ClO4−) anion in soil samples collected by the robotic arm (Hecht et al. 2009; Kounaves et al. 2010, 2014). Perchlorates are chemically inert at room temperature; however, if heated beyond a few hundred degrees, its four oxygen atoms are released, becoming very reactive oxidation vectors. It did not take long for investigators to recall that Viking had relied on thermal volatilization (TV; in other words heat) to release organics from soil samples (Navarro-González et al. 2010, 2011; Biemann and Bada 2011; Navarro-González and McKay 2011). If perchlorate had been also present in the soil at the two Viking lander locations, perhaps heating could explain the negative organic carbon results obtained? In fact, some simple chlorinated organic molecules (chloromethane and dichloromethane) had been detected by the Viking experiments (Biemann et al. 1977), but these compounds were interpreted to have resulted from a reaction between adsorbed residual methanol (a cleaning agent used to prepare the spacecraft) and hydrochloric acid (HCl). Today, the general consensus is that they were the outcome of heat-activated perchlorate dissociation and reaction with indigenous organic compounds (Steininger et al. 2012; Glavin et al. 2013; Quinn et al. 2013; Sephton et al. 2014; Goetz et al. 2016; Lasne et al. 2016).

This has been confirmed by measurements performed with the SAM (sample analysis at Mars) instrument on board Curiosity. The team detected oxygen (O2) released by the thermal decomposition of oxychlorine species [i.e., perchlorates and/or chlorates (Archer et al. 2016)], as well as chlorine-bearing hydrocarbons attributable to the reaction of oxychlorine species with organics compounds, both when they analysed modern sand deposits as well as when they drilled into much older rocks (Glavin et al. 2013; Freissinet et al. 2015). The exogenous delivery of meteoritic organics (abiotic) to the martian surface has been estimated at ~105 kg C/year, mostly in the form of polycyclic aromatic hydrocarbons (PAHs) and kerogen that may undergo successive oxidation reactions. Therefore, a meteoritic source could have contributed the organic precursors needed for producing the observed chlorobenzene and dichloroalkanes (Freissinet et al. 2015).

Summarizing, as a result of painstaking research performed over many years and involving multiple missions, we can conclude that:

  1. 1.

    There was plentiful liquid water on early Mars (at least during its first billion years). Since life seems to have appeared on our planet as soon as the environment allowed it, sometime between 4.4–3.8 Ga ago, it is likely that conditions also existed for the emergence of life on Mars, even if it was colder than Earth (Solomon 2005; McKay 2010; Strasdeit 2010; Yung et al. 2010);

  2. 2.

    Today, it is still possible to find organic molecules close to the surface, though most probably they have a meteoritic (not biogenic) origin;

  3. 3.

    Since the martian atmosphere is more tenuous than Earth’s, the ultraviolet (UV) radiation dose is higher than on our planet and will quickly damage exposed organisms or biomolecules; powerful oxidants exist that, when activated, can destroy the potential biosignatures we would like to study; and ionizing radiation penetrates into the uppermost meters of the planet’s subsurface, causing a slow degradation process that, operating over many millions of years, can alter organic molecules beyond the detection sensitivity of analytical instruments.

These are the very specific boundary conditions that Mars missions must contend with. Orbiters and landers exploring other worlds, for example the moons of Jupiter or Saturn, would need to adapt their search-for-life strategy to their particular environments’ geologic history—where are the potential biosignatures? how can we sample them? must we contend with radiation? But even if the details of the quest have to be tailored to the celestial body under investigation, it is possible to compile a somewhat universal list of observations having for objective to establish whether a location (on Mars or elsewhere) has hosted microbial life, past or present. In the work by Vago et al. (2017) we proposed such a list . We called this the ExoMars Biosignature Score because it is being developed while preparing for this mission. However, the list of biosignatures is quite exhaustive and encompasses more than those that the ExoMars rover will be able to assess (see Table 14.1).

Table 14.1 ExoMars Biosignature Score: A possible system to assign a confidence value (the score) to a group of robust observations aiming at establishing whether a location hosted life
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Please note that Table 14.1 does not include morphological changes with time (e.g. colony growth), movement (e.g. creature displacement), or experiments designed to trigger and observe active metabolic responses (as in Viking). These “more dynamic” expressions of possible present life would not be easy to verify.

In this chapter we address each of the observations in Table 14.1, explain what their positive verification would entail, provide examples (or sketches) to illustrate what the data could look like, and justify the need to provide evidence that several of the principal biosignatures have been demonstrated before positing that life detection has been achieved.

2 Geological Context and Biosignatures

For planetary surface missions, typically the characterisation of geological context begins early, with landing site selection, as investigators canvas interesting locations searching for those that best fit the mission’s scientific objectives.

Discovering candidate biosignatures embedded in a congruent geological landscape, that is, an environment that demonstrably possessed attributes conducive to the prosperity of microbial communities—for example, a long-lived, low-energy, aqueous or hydrothermal setting experiencing frequent fine sediment deposition—would help to increase substantially the confidence of any potential claim. For this reason we include “geological context” in the list of elements that bear investigation when searching for traces of life. However, some mission concepts do not afford this possibility. For example, a probe flying through an Enceladus plume seeking to trap and enrich organic molecule signatures would not have direct access to the geological setting where the organic molecules are being produced.

In general, microbial biosignatures can be grouped into three broad categories (Cady et al. 2003; Westall and Cavalazzi 2011) as follows: (1) cellular fossils that preserve organic remains of microbes and their extracellular matrices, as well their colonies and biofilms and mats, the study of which typically requires complex sample preparation and high-resolution instruments not currently available on landed space missions (Westall et al. 2011a); (2) bio-influenced fabrics and sedimentary structures (Noffke and Awramik 2013; Westall 2008, 2012; Davies et al. 2016), which provide a macroscale imprint of the presence of microbial biofilms that can be more readily identified, for example, laminated stromatolites; and (3) organic chemofossils preserved in the geological record (Parnell et al. 2007; Summons et al. 2008) that can be either primary biomolecules or diagenetically altered compounds known as biomarkers. Other, more tenuous or indirect information could complement the above, but we do not include them among the major biosignature categories. For example, metabolic effects preserved in the geologic record, such as mineral precipitation (e.g. magnetite by magnetotactic bacteria, Lefèvre and Bazilinski 2013, or carbonate, Dupraz et al. 2009), or leaching of elements in rocks (microbial corrosion, e.g. Foucher et al. 2010), or trace element concentration (e.g. Johannesson et al. 2014).

2.1 Morphological Biosignatures

The primordial types of microorganisms that could have existed on early Mars would have been prokaryotes similar to terrestrial chemotrophic microorganisms, i.e. of the order of a micron to a few microns in size. While the individual cells would be too small to distinguish, as on Earth, their permineralised or compressed microbial colonies and biofilms would be much larger.

In terrestrial marine (and other wet) environments, benthic microorganisms (e.g. those living in the seabed) form biofilms, highly organised microbial communities that are able to affect the accumulation of detrital sediments. Particle binding, bio-stabilisation, baffling, and trapping by biofilms can result in macroscopic edifices amenable to be recognised and studied with rover cameras and close-up imagers. In cases where sediment precipitation occurs in a repetitive manner, multi-layer constructions can ensue; for example, stromatolites constitute essential beacons of information, recording snapshots of microbial communities and environments throughout Earth’s history (Allwood et al. 2006, 2009, 2013). In particular phototrophs produce large amounts of extracellular polymeric substances in the biofilm . If the biofilm covers a large enough area experiencing similar conditions, often multiple organo-sedimentary structures can arise in regularly spaced groups. But the presence of microbes does not always lead to the emergence of noticeable macroscale bio-sedimentary formations. An example of a less conspicuous expression is the layering found in some typical early Earth volcanic lithic environments, where organisms have colonised the surface of ashfall particles, creating visible, carbon-rich, biofilms in various sediment horizons (Westall et al. 2011b, 2015).

Hereafter we present a few examples illustrating evidence of microbial colonisation in ancient Earth rocks (dated at 3.5–3.3 Ga). We chose to depict them at the scales at which they would be observed during a mission: panoramic (tens of meters to a few centimetres), close-up range (a few centimetres to several tens of microns), and microscopic (submicron). We include the latter for reference , although it is not possible today for robotic missions to prepare and study samples at such high magnification.

2.1.1 Panoramic Scale

Although the observation of identifiable biosignatures at the panoramic scale is not expected on Mars (unless fortuitous evolution led to the emergence of phototrophic organisms forming three-dimensional stromatolite-like structures), documentation of fine-scale layering in sediments deposited under water that may or may not be associated with accumulations of carbon can be made at this scale. Figure 14.1a shows finely layered sediments from the ~3.6 Ga Gale Crater whose structural features, such as cross-bedding and channel bedding, suggest deposition of sediments in a shallow water environment (cf. Grotzinger et al. 2014). Sedimentary structures in Fig. 14.1b from the 3.33 Ga Barberton Greenstone Belt, including ripple bedding, are indicative of deposition in a shallow water tidal environment (Westall et al. 2015).

Fig. 14.1
figure 1

(a) Panoramic view of an outcrop of the ~3.6 Ga Shaler bed (Yellowstone Bay, Gale Crater, Mars), showing finely layered volcanic sediments deposited in a lake. Source: Adapted from NASA/JPL-Caltech/MSSS. (b) Finely laminated volcanic sediments from the Josefsdal Chert (Barberton Greenstone Belt, 3.33 Ga) deposited in shallow coastal waters. An important difference between the two rocks is that those from the early Earth were silicified by hydrothermal fluids after deposition and are therefore harder and more resistant to erosion than those of Gale Crater. Source: F Westall. (c) Hydrothermal facies from Josefsdal showing chemical siliceous sediment (translucent layers) alternating with highly silicified, black layers. Source: F Westall

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The rock from Josefsdal in Fig. 14.1c is a hydrothermal chert. It includes translucent layers of chemically precipitated silica alternating with matt black layers that are faintly laminated (at the close-up scale). Observing sedimentary structures at this scale is important for interpreting the sedimentary context, but no further assertions can be made with respect to potential biosignatures.

2.1.2 Close-Up Scale

Close up observations are fundamental because they can provide more detailed information pertaining to the origin of the sediment particles as well as their mode of formation. For example, the size, sub-rounded habit and collision marks on the pebbles in Fig. 14.2a indicate strong aqueous transport (Edgar et al. 2017). Differential erosion of the fine, sandy component of the sediment has left the pebbles in relief. In comparison, Fig. 14.2b shows details of the Josefsdal Chert exposure shown in Fig. 14.1b exhibiting upwards grading of the sand, indicating passive settling of detrital particles deposited from airfall. Importantly, the close up images show that the black layers occur at the tops of the fining upwards cycles, indicating either that (1) they are composed of dark-coloured, hydrodynamically light detritus, and/or (2) the dark-coloured layers formed in situ on top of the sediment bedding planes, suggesting either diagenetic alteration of a precursor mineral into a dark phase (e.g. pyritisation) or the formation of microbial biofilms whose degraded organic remnants form dark-coloured kerogen.

Fig. 14.2
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(a) Close-up image of pebbles in a sandy matrix from the Shaler deposit at Yellowstone Bay, Gale Crater. Source: Adapted from Edgar et al. (2017). Surface textures on the pebbles and their sub-rounded aspect in a sandy matrix indicates subaqueous transport by relatively strong currents. (b) Close-up image of sedimentary textures, such as graded bedding (dotted arrow) in the finely laminated sediments in the Josefsdal Chert, Barberton, indicate airfall deposits into water. The black layers occur at the tops of the fining upwards sequences and could have an either detrital or in situ origin. Source: F Westall. (c) The hydrothermal sediments from the Josefsdal Chert are faintly laminated (arrow) but also show a mottled texture, e.g. bracketed area. Source: F Westall

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In contrast, the coarsely laminated hydrothermal sediments in the Josefsdal outcrop in Fig. 14.1c exhibit faint internal lamination in close-up imagery (Fig. 14.2c, arrow). What is of additional interest is the mottled texture exhibited by the matt black layers (N.B. the image has been lightened) indicating the presence of layers of small-scale internal clotted features some 100 s μm in size.

2.1.3 Microscopic Scale

The MAHLI-microscope image of sand grains in a dune in Gale Crater (Fig. 14.3a) show that good textural detail can be potentially obtained at the microscopic scale. However, in this specific instance, such detail is either not available or the dust cover on the surface of the rocks obscures potential detail. Microscopic observation, however, is essential for further interpretation, especially associated with a compositional method, such as Raman or IR spectroscopy. The examples in Fig. 14.3b–d show that some sample preparation is necessary.

Fig. 14.3
figure 3

(a) Sand grains on a dune in Gale Crater imaged by the instrument MAHLI. Individual grains are between 100–300 μm. Source: Adapted from Lapotre et al. (2017). (b) Optical microscope view in transmitted light of a sample from the Josefsdal Chert rock shown. Figure 14.2c showing that it contains particles 50–200 μm in size having a dark colour and producing a clotted texture . Source: courtesy F Foucher. (c) Optical detail of one particle (circled in b) with related Raman map (to the bottom) showing its mineralogy: anatase (blue), kerogen (green), silica (yellow–orange). Source: courtesy F Foucher. (d) SEM view of a colony of fossilised (silicified) microorganisms coating the surface of a similar volcanic grain from the 3.46 Ga Kitty’s Gap Chert, Pilbara. Source: Adapted from Westall et al. (2006)

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Thin sections (30 μm thick) of the rock in Fig. 14.2c were made and investigated under optical microscopy in transmitted light and Raman spectral mapping. The high spatial resolution study of details of the clotted texture (which were just faintly visible in close-up imagery) showed that the rock consists of volcanic particles largely coated with a black phase, which could be either kerogen or an oxide. The Raman mapping documented alteration of the underlying volcanic particle to anatase, a process that takes place in the presence of water. It also showed that the particle was thoroughly coated by kerogen. The irregular thickness of the coating (Westall et al. 2015) suggests in situ growth, such as colonisation by chemotrophic microbes, rather than coating by abiogenic carbon, which would produce a conformable layer.

In Fig. 14.3c, the rock surface was treated with hydrofluoric acid (HF) to highlight the differential mineralogy of the different components within the rock, thus preferentially corroding the silica matrix of the sediment and exposing the “dirty”, carbon-rich components (i.e. carbonaceous microfossils). Visualisation of the <1 μm-sized microfossils is only possible using high powered electron microscopy. Even with this kind of compositional and textural data, interpretation of such features can be controversial.

2.2 Chemical Biosignatures

When considering molecular biosignatures , the first obvious set of targets is the ensemble of primary biomolecules associated with active microorganisms, such as amino acids, proteins, nucleic acids, carbohydrates, some pigments, and intermediary metabolites. Detecting the presence of these compounds in high abundance would be diagnostic of extant life, but unfortunately on Earth they degrade quickly once microbes die.

Lipids and other structural biopolymers, however, are biologically essential components (e.g. of cell membranes) known to be stable for billions of years when buried (up to 1.64 Ga on Earth, Brocks et al. 2005). It is the recalcitrant hydrocarbon backbone that is responsible for the high-preservation potential of lipid-derived biomarkers relative to that of other biomolecules (Eigenbrode 2008).

Along the path from primary compound to molecular fossil, all biological materials undergo in situ chemical reactions dictated by the circumstances of transport, deposition, entombment, and post-depositional effects on the original organisms. The end product of diagenesis is macromolecular organic matter, which, through the loss of superficial hydrophilic functional groups, slowly degrades into the solvent-insoluble form of fossil carbonaceous matter called kerogen, as seen in the example from the Josefsdal Chert (Fig. 14.3c), The heterogeneous chemical structure of the kerogen matrix can preserve patterns and distribution diagnostic of biosynthetic pathways. Kerogen also possesses molecular sieve properties allowing it to retain diagenetically altered biomolecules (Tissot and Welte 1984).

Apart from the direct recognition of biomolecules and/or their degradation products (which is not illustrated here), other characteristics of bioorganic compounds include the following (Summons et al. 2008, 2011).

2.2.1 Enantiomeric Excess

In the case of chiral molecules (those that can exist in either of two non-identical mirror image structures known as enantiomers), life forms synthesise exclusively one enantiomer, for example, left-handed amino acids (l-amino acids) to build proteins and right-handed ribose (d-ribose) for sugars and the sugars within ribonucleic acid (RNA) and deoxyribonucleic acid (DNA). Opposite enantiomers (d-amino acids and l-ribose) are neither utilized in proteins nor in the genetic material RNA and DNA.

The use of pure chiral building blocks is considered a general molecular property of life. For example, gradual loss of enantiomeric excess in amino acids, i.e. racemisation, in fossil shells has been used for dating Quaternary fossils (called aminochronology, Wehmiller 1993).

2.2.2 Molecular Weight Clustering of Organic Compounds

Many important biochemicals exist in discrete molecular weight ranges (e.g. C14–C20 lipid fatty acids). For this reason, the molecular weight distribution of biologically derived matter exhibits clustering; it is concentrated in discrete clumps corresponding to the various life-specialized families of molecules (Summons et al. 2008). This is in contrast to the molecular weight distribution for cosmic organics (Ehrenfreund and Charnley 2000; Ehrenfreund and Cami 2010): the relative abundance for abiotic volatiles is uniform and drops off as the carbon number increases. For example, ToF-SIMS analysis of well-preserved, silicified kerogen from the 3.33 Ga Josefsdal Chert in South Africa, show just such clustering of the organic molecules (Westall et al. 2011b).

2.2.3 Repeating Constitutional Subunits

Many biological products (e.g. proteins and nucleic acids) are synthesized from a limited number of simpler units. This can leave an identifiable molecular weight signature even in fragments recovered from highly derived products, such as petroleum. For example, in the case of material containing fossil lipids we would expect to find a predominance of even-carbon numbered fatty acids (C14, C16, C18, C20). This is because the enzymes synthesizing fatty acids attach two carbon atoms at a time (in C2H4 subunits) to the growing chain. Other classes of biomolecules can also exhibit characteristic carbon chain-length patterns, for example, C15, C20, C25 for acyclic isoprenoids constructed using repeating C5H10 blocks.

This phenomenon is illustrated by analysis of kerogen from the Bitter Spring Formation (850 My) in northern Australia where the aromatic molecules linked by short-chain n-alkyl moieties exhibit odd carbon number predominance, perhaps indicative of bacterial cell wall lipids (Imbus and McKirdy 1993).

2.2.4 Systematic Isotopic Ordering at Molecular (Group) Level

Biological molecule building blocks , in particular some functional groups, can show significant differences in their degree of 13C incorporation relative to 12C. The “repeating subunit” conformation of biomolecules can result in an observable isotopic ordering in the molecular fingerprint.

2.2.5 Bulk Isotopic Fractionation

The isotopic fractionation of stable elements such as C, H, O, N, S, and Fe can be used as a signature to recognize the action of biological pathways. Although the qualitative chemical behaviour of the light and the heavy isotope is similar, the difference in mass can result in dissimilar bond strength and reaction rates. Thus, the isotopic discrimination associated with organic biosynthesis (which alters the natural equilibrium between C isotopes in favour of the lighter variant) is principally responsible for determining the 13C/12C ratios in terrestrial organic and inorganic crustal reservoirs.

Although interesting, we do not consider bulk isotopic fractionation a robust biosignature when applied to locations or epochs for which we have scant knowledge of sources and sinks. In the specific case of carbon, 13C/12C ratios may serve as reliable biosignatures for past or present life only if the key components of the C-cycling system (applicable at the time of deposition and since then) are well constrained (Summons et al. 2011). This is certainly not the case for Mars (or other bodies), and one can also wonder to what extent we are sure about our own past carbon dynamics when analysing very ancient samples .

Despite the above reservations we are willing to include bulk isotopic fractionation in this list, but with the caveat that it should be used in association with other, less indirect, biosignatures.

On Earth, groundbreaking work on isotopic signatures and their preservation in ancient rocks was made by Manfred Schidlowski (1988) and ever since, the δ13C signature is considered as a useful accompanying biosignature. Schidlowski (1988) documented isotopically light carbon in 3.8 Ga rocks containing carbon from the Isua Greenstone Belt in Greenland, although Westall and Folk (2003) noted the presence of recent (<8000 year-old) fossilised endolithic microorganisms in the rocks analysed. More recent studies of Isua rocks using the graphite combustion method do indeed confirm the presence of isotopically light carbon with δ13C values between −10 and −20‰ (Ohtomo et al. 2014).

2.2.6 Preservation of Organic Matter

Effective chemical identification of biosignatures requires access to well-preserved organic molecules.

Because the martian atmosphere is more tenuous than Earth’s, three important physical agents reach the surface of Mars (and other airless Solar System bodies of interest) with adverse effects for the long-term preservation of biomarkers: (1) The UV radiation dose is higher than on our planet and will quickly damage exposed organisms or biomolecules. (2) UV-induced photochemistry is responsible for the production of reactive oxidant species that, when activated, can also destroy chemical biosignatures. The diffusion of oxidants into the subsurface is not well characterized and constitutes an important measurement that future missions must perform. Finally, (3) ionizing radiation penetrates into the uppermost meters of the planet’s subsurface. This causes a slow degradation process that, operating over many millions of years, can alter organic molecules beyond the detection sensitivity of analytical instruments. Radiation effects are depth dependent: the material closer to the surface is exposed to higher doses than that buried deeper. Therefore, the molecular record of ancient martian life, if it ever existed, is likely to have escaped radiation and chemical damage only if trapped in the subsurface for long periods. Studies suggest that a subsurface penetration in the range of 2 m is necessary to recover well-preserved organics from the very early history of Mars (Kminek and Bada 2006), assuming there has been some help from additional, recently eroded overburden (Dartnell et al. 2007, 2012; Parnell et al. 2007; Pavlov et al. 2012).

It is also important to notice that large diurnal temperature excursions are typical for most bodies lacking a dense atmosphere able to provide much thermal inertia. For example, at midlatitudes on Mars, noon surface temperatures hover around 0 °C, but quickly plummet to −120 °C during the night. This can work to our advantage. Just 50 cm underground the temperature remains close to the average between day and night, some −60 °C. Thus, the martian subsurface effectively behaves like a freezer, slowing down the chemical reactivity of damaging oxidants, greatly contributing to the chemical preservation of organic molecules. The weaker gravity also means that diagenetic processes that would normally affect the preservation of molecules are comparably weaker on Mars. The same column of rock produces less pressure than on a higher gravity planet. It is therefore likely that fragile molecules that would quickly degrade on Earth could last much, much longer on Mars preserved sufficiently deep in the cold subsurface.

3 Conclusion

We have proposed a possible scoring system for assigning a confidence value to a group of observations aiming to establish whether a location hosted microbial life.

We find that there is value in defining a set of measurements and rules to guide our preparations and help with the interpretation of any findings once we have reached our destination, Mars or elsewhere.

A closer examination of Table 14.1 reveals that, if we could tick all possible biosignatures, assigning maximum points with a perfect chemical background, the score would be 200. However, in the work by Vago et al. (2017), we claimed we only need a value of 100 to establish that there was/is life. This indicates that it is not necessary to verify all possible biosignatures, but that it is mandatory to provide evidence that a few of the principal biogenicity criteria or indicators are indeed demonstrated. Chemical biosignatures are awarded a higher importance because they provide “more direct” evidence of biogenicity than the other categories for which bioinfluence is “inferred”. Note, however, that in certain circumstances morphological biosignatures can be equally well-constrained.

Life-seeking missions to other planets should target as many biosignatures indicated in Table 14.1 as possible. We claim that their discoveries will not be conclusive unless such missions include powerful analytical chemistry capabilities that can allow for the unambiguous identification of key biosignatures of biomolecules and their degradation products.