The same high‐radiation and oxidizing conditions that limit the habitability of the Martian surface today would also destroy exposed organic remains. The average galactic cosmic ray dose rate measured in Gale Crater was ~0.2 mGy/day (Hassler et al., 2014 ); energetic particles associated with solar flares and coronal mass ejections can occasionally produce dose rates orders of magnitude higher, and the UV flux combined with strong oxidants also destroys organic matter (e.g., ten Kate et al., 2005 ; Wadsworth & Cockell, 2017 ). Although the exact nature of the early Martian climate is still uncertain, the Noachian‐Early Hesperian radiation flux was attenuated by a thicker atmosphere that supported stable liquid water at the surface (Grotzinger et al., 2014 ; Mahaffy et al., 2015 ), and potentially also by an Earth‐like magnetic field, although this may have been lost by the early Noachian (e.g., Lillis et al., 2013 ). The strong oxidants presently in Martian soil may have been less concentrated or absent in the wetter conditions and milder radiation environment, although chlorine isotope measurements by the Curiosity rover in Gale Crater suggest perchlorate production during the Hesperian (Carrier & Kounaves, 2015 ; Farley et al., 2016 ). Any organisms dwelling on early Mars would therefore presumably have been degraded primarily as a result of biological heterotrophy and enzymatic autolysis, as on Earth. Although low temperatures and a lack of burrowing or grazing animals would have inhibited decay and favored preservation, any fossil record on Mars would be biased toward robust, decay‐resistant biological materials, such as biominerals/organominerals (Perry et al., 2007 ) or thick microbial sheaths, and toward environments that promoted preservation. Favorable circumstances for preservation would have included rapid burial, a high accumulation rate of organic remains, the presence of mineralizing fluids, and the inhibition of decay. Research on the nature, distribution, and quality of fossils representing both microsoft‐ and macrosoft‐bodied organisms on Earth, supported by decay experiments on cells and tissues and the environmental conditions required to preserve them (“experimental taphonomy”; Briggs & McMahon, 2016 ), provides essential guidelines to the minerals, lithologies, and facies most likely to host microbial fossils on Mars (Table 1 ).

It is thus highly likely that various Martian paleoenvironments were habitable, and plausible that they were inhabited, but could they have preserved fossils? On Earth, most organisms fail to fossilize because their remains are physically destroyed, chemically oxidized or dissolved, digested by their own enzymes, or consumed by other organisms. Fossilization only occurs when processes of preservation outpace degradation (Allison, 1988 ). Fossilization usually begins when organisms are buried in sediment or entombed in minerals, fixing their remains in place, and reducing their exposure to degradative processes including heterotrophic consumption (animal, fungal, or bacterial). Minerals may replicate the morphology of the buried organism, cement the sediment around it, or fill the remaining void following decay. Microbes commonly self‐fossilize by entombing themselves in the mineral byproducts of their own metabolic activity (see section 4 ). Organic cells and tissues are usually short lived, but they can serve as a template for the nucleation of more permanent minerals or their precursors, which in turn can stabilize the original organic matter, particularly when clay minerals form (Gabbott, 1998 ; Naimark, Kalinina, Shokurov, Markov, et al., 2016 ). Organic material can also survive when passively entombed by rapid precipitation, as in Precambrian marine cherts and Phanerozoic hydrothermal sinters (e.g., Campbell et al., 2015 ; Pancost et al., 2006 ; Schultze‐Lam et al., 1995 ; Yee et al., 2003 ). Organic molecules with robust chemical backbones can be indicative of general or specific biological origins (i.e., “biomarkers” or “molecular fossils”) and are common in fine‐grained sedimentary rocks (Peters et al., 2005 ; Summons & Lincoln, 2012 ).

The Martian surface has been cold and predominantly dry for at least the last three billion years (i.e., the Amazonian Period, immediately following the Hesperian), but the subsurface could have sustained stable reservoirs of geothermally heated liquid water for much of this time, representing a long‐lived habitat that could have exchanged living cells with shorter‐lived habitats at the surface (Boston et al., 1992 ; Clifford et al., 2010 ; Ehlmann et al., 2011 ; Fisk & Giovannoni, 1999 ; Thomas et al., 2017 ; Travis et al., 2003 ). Transport from the surface would probably have been required to provide the initial inoculum or at least chemical precursors to life (e.g., Patel et al., 2015 ). Once established, however, a “deep biosphere” would have been protected from the deteriorating conditions at the surface and able to persist far longer given indigenous sources of energy and nutrients (Des Marais, 2010 ; Westall, Foucher, et al., 2015 ). Potentially analogous subterranean ecosystems have been studied most intensively in the Witwatersrand Basin of South Africa, where anaerobic microbial populations achieve densities of at least 10 3 –10 4 cells/mL at depths of 3–4 km (Lin et al., 2006 ; Moser et al., 2005 ). Life in this setting appears to have been isolated from the surface for millions of years (Lin et al., 2006 ) and relies largely on abiotic carbon sources and electron donors, the latter dominated by molecular hydrogen supplied by water‐rock reactions (Lin et al., 2005 ; Sherwood Lollar et al., 2006 ). There is clear evidence that such reactions occurred on Mars, including the presence of the mineral serpentine at the Martian surface and in Martian meteorites (Blamey et al., 2015 ; Changela & Bridges, 2011 ; Ehlmann et al., 2010 ). The putative detection of intermittent traces of methane in the Martian atmosphere—which could be a product of water‐rock reactions or even of hydrogen‐oxidizing microbes themselves—may imply the persistence of a habitable environment into geologically recent time (Krasnopolsky et al., 2004 ; Mumma et al., 2009 ; Webster et al., 2015 ).

There is no compelling evidence for life on Mars, now or in the geological past. However, there is now a very strong case that the surface of early Mars was habitable. Orbital data reveal fluvial valley networks draining thousands of square kilometers, exhumed meandering and branched distributary channels, paleolake deposits in topographic depressions, and alluvial fans/deltas entering these lakes, all of which reflect sustained precipitation and subaqueous sediment transport during the Noachian and Hesperian periods (Cabrol & Grin, 1999 ; Fassett & Head, 2008 ; Goudge et al., 2016 ; Grant et al., 2008 ; Grotzinger, Gupta, et al., 2015 ; Malin & Edgett, 2003 ; Metz et al., 2009 ; Moore & Howard, 2005 ). Some fan‐shaped deposits, possibly deltaic, have been interpreted as aligned along the shoreline of a large ocean (Di Achille & Hynek, 2010 ; DiBiase et al., 2013 ) that would have covered the northern lowlands, although this is controversial. At Gale Crater, however, the Mars Science Laboratory mission (Curiosity rover) has explored the sedimentary record of a Late Noachian/Early Hesperian paleolake that persisted for thousands to millions of years, with evidence for mild salinity, moderate pH, and local redox gradients (Grotzinger et al., 2014 ; Grotzinger, Gupta, et al., 2015 ; Hurowitz et al., 2017 ). The presence of liquid water at the surface over this length of time implies a denser atmosphere, which should also have minimized the radiation flux. Tantalizingly, these rocks contain organic molecules (Freissinet et al., 2015 ) preserved under relatively reducing conditions (Grotzinger et al., 2014 ). Such Noachian‐Hesperian water bodies could have hosted microbial life, which was present on Earth during this time interval (Grotzinger et al., 2014 ). There is experimental support for the long‐standing idea that meteorites could have transported viable microbial cells between the two planets, although the ecological obstacles to continued growth on arrival may be considerable (Fajardo‐Cavazos et al., 2005 ; Foucher et al., 2010 ; Mileikowsky et al., 2000 ; Shuster & Weiss, 2005 ; Weiss et al., 2000 ).

The search for evidence of life on Mars is one of the outstanding scientific challenges of our time. Low temperatures and pressures, intense ionizing radiation, oxidizing soil chemistry, and low levels of thermodynamic water activity greatly reduce the possibility of life at the Martian surface today. However, extensive ancient valley networks and sedimentary rocks laid down in the Noachian and Hesperian Periods of Martian history (respectively, ~4.0–3.6 billion years ago [Ga] and ~3.6–3.0 Ga; Fassett & Head, 2011 ; Werner & Tanaka, 2011 ) are suggestive of much more clement global conditions, including widespread liquid water at the surface. The uncertain timing and duration of water availability are still debated and have important implications for the potential viability, complexity, distribution, and preservation potential of any life that arose on early Mars. Nevertheless, Noachian and Hesperian rocks have long been recommended as a target for fossils, that is, the physical or chemical remains of organisms and their activities (e.g., McKay & Stoker, 1989 ). On Earth, these remains are preserved as casts or molds in sediment, as mineral coatings or replacements, and as surviving biominerals, organic matter, or stable isotopic signatures. Locating samples on Mars that are likely to preserve such signals and studying them with appropriate instrumentation, both in situ and possibly after returning them robotically to Earth, are now strategic priorities for both NASA (National Aeronautics and Space Administration) and ESA (the European Space Agency). One of the primary objectives of NASA's Mars 2020 rover is to cache a variety of drill‐core rock samples from astrobiologically promising paleoenvironments; some of these samples may be returned to Earth by subsequent missions for analysis (Farley & Williford, 2017 ; Mustard et al., 2013 ). The ESA‐Roscosmos ExoMars 2020 rover will also be equipped to test technologies required for subsequent sample‐return missions (Vago et al., 2017 ). Understanding how, why, and where morphological, molecular, or isotopic biosignatures might have survived on Mars will significantly enhance the success of these missions by informing the selection of landing sites, rover traverse pathways, and sampling strategies.

2 Potentially Fossiliferous Rocks and Minerals on Mars

2.1 Secondary Minerals in Basalt The likelihood that the surface of Mars has been dry for the last several billion years has directed attention to the habitability of fractures and pores in deep basaltic rocks and sediments, where liquid water may have been sustained by geothermal heat (e.g., Cockell, 2014a, 2014b; Hays et al., 2017; Michalski et al., 2013). However, the potential for preserving fossils within the secondary minerals that form in such rock‐hosted habitats remains unclear. Microscopic textures in basalt, ranging from simple filled or unfilled pits to elaborate helical “tunnels,” are widely reported as “bio‐alteration” features or microbial “trace fossils” (e.g., McLoughlin et al., 2009; Staudigel et al., 2008; Thorseth, 2011), but in most cases an abiotic origin cannot be excluded (Fisk et al., 2013). More compelling are reports of filamentous and rod‐shaped structures mineralized in basalt void space on Earth by low‐temperature submarine or subterranean hydrothermal activity (e.g., Cavalazzi et al., 2011; Hofmann & Farmer, 2000; Ivarsson et al., 2016; McKinley et al., 2000; Peckmann et al., 2008; Schumann et al., 2004). Such structures range from micrometer to millimeter in scale and are typically preserved three‐dimensionally in clay, iron oxides, and other minerals within open‐ or mineral‐filled vesicles and fractures. Some appear to have been mineralized in anoxic conditions (e.g., Bengtson et al., 2017; Ivarsson et al., 2016; Schumann et al., 2004). Unfortunately, it has not been demonstrated that any of these structures represent Mars‐relevant chemoautotrophs. Indeed, subsurface productivity on Earth is strongly dependent on the heterotrophic remineralization of carbon originally fixed at the surface by photosynthesis (Kallmeyer et al., 2012; McMahon & Parnell, 2014). As such, the bulk of Earth's massive deep biosphere, and presumably also its fossil record, is a poor analog for any ancient or modern Martian equivalent which, in the absence of a productive surface biosphere, would be much smaller and dominated by chemoautotrophs, not heterotrophs. Lithoautotrophic communities in deep groundwaters in South Africa and the Canadian Shield are more relevant to scenarios for life on Mars (e.g., Moser et al., 2005; Onstott et al., 2009). However, these stagnant aquifers have chemically equilibrated with their mineralogical environment following long isolation from the surface (Lippmann et al., 2003; Holland et al., 2013) and are therefore unlikely to mineralize actively except during rare tectonic events (McNutt et al., 1990; Sherwood Lollar et al., 2007). To our knowledge, there are no reports of mineralized cells found at these sites nor any definitive fossil record from elsewhere of ancient microbial ecosystems analogous to these. Mineralized fractures and void spaces are difficult to resolve and interpret at the spatial scale of orbital data from Mars, and even on Earth probably do not usually yield robust indigenous biosignatures. A better understanding of the distribution, degradation, and preservation of Mars‐relevant microbes in Mars‐relevant subsurface environments is therefore needed before landing sites can be selected to search for a fossil deep biosphere on Mars.

2.2 Evaporite Salts Evaporite salts, including Ca/Mg/Fe/Al‐sulfates and chlorides, are widespread in basins on Mars, commonly occurring in Noachian and Hesperian‐aged terrains (Ehlmann et al., 2016; Gendrin et al., 2005; Glotch et al., 2010; Murchie et al., 2009; Osterloo et al., 2008; Wray et al., 2011). In some cases, these salts are associated with paleoenvironments inferred to be playa like or lacustrine with strong potential for habitability (Ehlmann & Edwards, 2014; Grotzinger et al., 2005), although salinities may have been so high in Mg‐sulfate brines as to render them uninhabitable (Knoll & Grotzinger, 2006; Tosca et al., 2008). Such evaporite minerals have the potential to trap organisms and preserve them as organic fossils which, like those in chert (see below), are readily visible in transmitted light. Gypsum‐permineralized carbonaceous microfossils interpreted as algae and bacteria, the latter comparable to anaerobic nitrate‐reducing sulfide oxidizers, have recently been found in Permian, Miocene, and recent evaporites (Pierre et al., 2015; Schopf et al., 2012). Calcified organic‐rich microbial filaments have also been found in Miocene gypsiferous stromatolites (Allwood et al., 2013; see section 2.3 below for discussion of stromatolites). It has been suggested that fluid inclusions in halite may preserve viable microbial cells over ~100 Myr, although this is controversial because of well‐founded doubts that such cells could repair DNA damage induced by natural radiation over these timescales (Fish et al., 2002; Jaakkola et al., 2016; Satterfield et al., 2005; Vreeland et al., 2000). More research needs to be done on potential controls on the presence and long‐term persistence of morphological and molecular fossils in evaporites before prioritizing these minerals in landing‐site selection. However, vertical and lateral patterns in sulfates and other salts in lacustrine environments can record important information about paleoenvironmental and physiochemical conditions (Bristow & Milliken, 2011). Thus, lacustrine evaporites would be a target for sampling if encountered in the course of a rover traverse, especially if associated with other evidence for habitable water chemistry. Both Curiosity and Opportunity have also encountered sulfate veins in postburial fractures (Caswell & Milliken, 2017; Vaniman et al., 2014), representing diagenetic precipitates from briny groundwater rather than bottom‐nucleated growth at the sediment‐water interface, the common style of evaporitic deposition on Earth. The potential for such vein sulfates to yield biosignatures is not well understood.

2.3 Phosphate and Carbonate Laminated and nodular phosphates, and to a lesser extent carbonates, are important sources of well‐preserved fossil microbes and organic matter on Earth (e.g., Figure 1a; Knoll et al., 1993; Morais et al., 2017; Xiao et al., 2014). Unfortunately, similar deposits have not yet been identified on Mars. Small amounts of phosphorus have been detected at Gale Crater, both as igneous‐sourced detrital crystalline apatite in sandstone and as a component of an amorphous/poorly crystalline phase in mudstone, which may be secondary (Forni et al., 2015; Rampe et al., 2017). Magmatic and metasomatic phosphate also occur in Martian meteorites, as do trace amounts of carbonate. Carbonate in the soil at the Phoenix landing site is considered to have formed in situ recently from CO 2 dissolved in thin water films (Boynton et al., 2009); it is absent or below detection limits at Gale Crater (Bristow et al., 2017). Climate models indicate that temperatures would have been too low under the “faint young Sun” to sustain large volumes of water on early Mars without abundant atmospheric CO 2 , conditions that would have promoted widespread carbonate mineralization, particularly as basalt would have buffered pH (Niles et al., 2013; Wordsworth, 2016). Thus, the lack of bedded carbonate on the Martian surface is difficult to reconcile with the abundant evidence for wet conditions in the Noachian‐Hesperian. The early Martian atmosphere may have been warmed by other greenhouse gases with minimal contribution from CO 2 and hence insignificant carbonate formation (e.g., Bristow et al., 2017). Alternatively, large carbonate deposits may be concealed beneath alteration assemblages, lava flows, or soil (Clark, 1999) and have yet to be discovered; indeed, isolated carbonates formerly buried to several kilometers have been detected in craters (Michalski & Niles, 2010; Wray et al., 2016). However, the origin of these carbonates is not clear, and impact metamorphism might have damaged biosignatures in such exposures. Figure 1 Open in figure viewer PowerPoint Girvanella) in limestone, upper Cambrian Campbell's Member, western Newfoundland. Image courtesy of S. Pruss, Smith College. (b) Stromatolites in chert, Archean Strelley Pool Formation, Western Australia. (c) Stromatolites in limestone, Paleoproterozoic Rocknest Formation, Wopmay Orogen, northwest Canada. (d) Stromatolites in sandstone, Neoproterozoic Witvlei Group, Namibia. (e) Filamentous and coccoidal microfossils in chert, Paleoproterozoic Gunflint Formation, Ontario, Canada. Image courtesy of A. H. Knoll, Harvard University. (f) Mat‐forming colonial coccoidal cyanobacteria in chert, Neoproterozoic Min'yar Formation. (g) Wrinkle structures in siltstone draped over conglomerate, middle Cambrian March Point Formation, western Newfoundland. Image courtesy of S. Pruss, Smith College. (h) Organically preserved cyanobacteria (Symplassosphaeridium sp.) macerated from shale, upper Mesoproterozoic Iqqittuq Formation, Arctic Canada. Image courtesy of H. Agić, University of California, Santa Barbara. Scale bar: (a) 200 μm, (e) 75 μm, (f) 625 μm, (g) 60 mm, and (h) 120 μm. The scale of Figures Terrestrial fossils that inform the search for life on Mars: (a) Calcified cyanobacterial sheaths () in limestone, upper Cambrian Campbell's Member, western Newfoundland. Image courtesy of S. Pruss, Smith College. (b) Stromatolites in chert, Archean Strelley Pool Formation, Western Australia. (c) Stromatolites in limestone, Paleoproterozoic Rocknest Formation, Wopmay Orogen, northwest Canada. (d) Stromatolites in sandstone, Neoproterozoic Witvlei Group, Namibia. (e) Filamentous and coccoidal microfossils in chert, Paleoproterozoic Gunflint Formation, Ontario, Canada. Image courtesy of A. H. Knoll, Harvard University. (f) Mat‐forming colonial coccoidal cyanobacteria in chert, Neoproterozoic Min'yar Formation. (g) Wrinkle structures in siltstone draped over conglomerate, middle Cambrian March Point Formation, western Newfoundland. Image courtesy of S. Pruss, Smith College. (h) Organically preserved cyanobacteria (sp.) macerated from shale, upper Mesoproterozoic Iqqittuq Formation, Arctic Canada. Image courtesy of H. Agić, University of California, Santa Barbara. Scale bar: (a) 200 μm, (e) 75 μm, (f) 625 μm, (g) 60 mm, and (h) 120 μm. The scale of Figures 1 c and 1 d is indicated by a Swiss army knife, hammer, and lens cap, respectively. Despite the apparent lack of bedded carbonate on Mars, carbonates formed at low temperatures (~18°C) are present in the ~4.1 Ga Martian meteorite ALH84001 (Halevy et al., 2011). In addition, carbonates of possible hydrothermal origin offer an alternative target for biosignature detection. The Mars Reconnaissance Orbiter identified magnesium carbonate associated with olivine and clays in the Nili Fossae region (Ehlmann, Mustard, Murchie, et al., 2008), and Spirit discovered carbonate‐rich (16–34 wt %) outcrops (named the Comanche outcrops) of similar composition in Gusev Crater (Morris et al., 2010). These carbonates probably formed through the aqueous alteration of mafic precursors by hydrothermal activity. The evidence for hydrothermal activity in Gusev Crater may indicate a genetic similarity between the carbonates there and volcanism‐related, nonmarine, Mg‐rich travertines on Earth. Some young travertines yield organic biomarkers (e.g., Jorge‐Villar et al., 2007) and microbial microfabrics (Riding, 1991). Submarine carbonate vent chimneys can likewise preserve molecular fossils as well as isotopic biosignatures (e.g., Brazelton et al., 2006; Lincoln et al., 2013; Méhay et al., 2013;). Molecular, microfossil, and isotopic biosignatures in carbonates are vulnerable to damage by fluid throughflow, chemical alteration, and recrystallization over geological time. Young hydrothermal carbonates contain cellular and molecular fossils (e.g., Zhang et al., 2004), and cellular preservation by iron and carbonate minerals has been reported from Jurassic travertines where Ostwald ripening of calcite seems to have inhibited diagenetic alteration (Potter‐McIntyre et al., 2017). Precambrian travertines lack such biosignatures, which may reflect sustained alteration processes on Earth that would be less severe on Mars (Brasier et al., 2013; see section 5 below). However, these rocks do commonly contain stromatolites, that is, layered conical, domal, columnar, or branching macroscopic growth structures attached to a surface and formed by carbonate precipitation and/or the trapping and binding of sediment (Figures 1b–1d; Bosak et al., 2013; Grotzinger & Knoll, 1999; Riding, 1999). Microbes are commonly implicated in these processes, but it has long been clear that not all stromatolite‐like features are necessarily biological, especially those formed by precipitation (rather than trapping and binding). This complicates the interpretation of Precambrian precipitated stromatolites and those that have undergone substantial diagenesis (Allwood et al., 2009; Grotzinger & Knoll, 1999; Grotzinger & Rothman, 1996). Triangular structures exposed perpendicular to bedding on a weathered, heavily metamorphosed carbonate in the Isua Supracrustal Belt in Greenland, for example, which were interpreted by Nutman et al. (2016) as Earth's earliest stromatolites, are morphologically ambiguous (their 3‐D structure is unreported) and lack organic carbon or other evidence to confirm biogenicity. Although microfossils are rare in carbonate stromatolites, studies of Precambrian examples and modern analogs have identified structures and morphologies with a high potential to record biological activity (e.g., Allwood et al., 2006; Beukes & Lowe, 1989; Bosak et al., 2009, 2010; Dupraz et al., 2004; Grey, 1994; Hoffman, 1976; Jones et al., 1997, 1998; Komar et al., 1965; Reid et al., 2000; Sim et al., 2012; Sumner, 1997). Only recently, however, through a combination of theory, experiment, and field observations, have we begun to understand the processes that produce robust morphological biosignatures in macroscopic stromatolite‐like structures as old as three billion years (Batchelor et al., 2000; Batchelor et al., 2004; Batchelor et al., 2005; Bosak et al., 2013; Cuerno et al., 2012; Dupraz et al., 2006; Mariotti, Perron, et al., 2014, Mariotti, Pruss, et al., 2014; Petroff et al., 2010, 2013; Sim et al., 2012; Walter et al., 1976) or in microscopic textures (Bosak et al., 2009; Bosak et al., 2010; Bosak et al., 2013; Mata et al., 2012). Although most stromatolites are too small to be identified remotely, they would be readily observable by a rover on Mars and would be a prime target for astrobiological sampling. More generally, however, further research is needed to clarify the potential for biosignature preservation in carbonates similar to those so far encountered on Mars.

2.4 Hydrothermal Silica Hydrothermal systems, both at and below the paleosurface, have long been recognized as likely habitable sites with the potential to preserve fossils (e.g., Farmer & Des Marais, 1999; McKinley et al., 2000; Walter & Des Marais, 1993). Some Noachian terrains are inferred to record mineral alteration by hydrothermal fluids that passed through the Martian upper crust prior to excavation by impact cratering (Ehlmann et al., 2009, 2011; Michalski et al., 2013). The thermal afterglow of impacts themselves can drive hydrothermal circulation in the vicinity of craters (Osinski et al., 2013), which may produce postimpact silica and sulfate veins, as well as Al‐rich clays (e.g., as revealed in Endeavour Crater by the Opportunity rover; Arvidson et al., 2014). A recent global survey of crater central peaks using Mars Reconnaissance Orbiter data has shown that ~22% of those with hydrated minerals show spectral evidence for hydrated (opaline) silica associated with uplifted materials, possible impact melt deposits, and various unconsolidated materials (Sun & Milliken, 2015). Silica dissolved and mobilized at depth stays in solution at high temperatures, but the expression of hydrothermal systems at the cool sediment‐water or sediment‐atmosphere interface induces rapid, massive surface precipitation. The resulting sinter deposits typically preserve microbial filaments as silicified casts, molds, and coatings, often in such density and abundance that they constitute a large fraction of the rock and determine its macroscopic texture (e.g., Cady & Farmer, 1996; Munoz‐Saez et al., 2016; Trewin et al., 2003). Young examples yield a wide range of lipid biomarkers representative of hot spring organisms (e.g., Gibson et al., 2008; Kaur et al., 2008; Pancost et al., 2006). Other biosignatures can include fenestrae representing silicified bubbles of microbially generated gases (Bosak et al., 2009, 2010; Mata et al., 2012) and millimeter‐scale laminated fabrics arising from the interplay of biofilms and silicifying fluids (e.g., Konhauser et al., 2004). Opaline siliceous sinter transforms to solid microcrystalline or cryptocrystalline forms (cristobalite, tridymite, and quartz, i.e., chert) during burial, which may preserve the organic remains of eukaryotes and prokaryotes at submicron resolution. The best‐known example is the Devonian (~410 Ma) Rhynie Chert in Scotland, which preserves plants, animals, fungi, and bacteria entombed and permineralized with silica (Trewin, 1993, 1996) and yields well‐preserved organic biomarkers (e.g., Preston & Genge, 2010; Qu et al., 2015). Finely laminated Archean cherts containing hydrothermally silicified biofilms also preserve some organic matter, but the high metamorphic grades of these rocks (Westall, Campbell, et al., 2015) ensure that any molecular biosignatures have been erased. Curie point pyrolysis of an Archean chert from the Warrawoona Formation of the Pilbara Craton in Western Australia yielded alkane/alkene doublets with a slight odd/even preference (Derenne et al., 2008), an indisputable biosignature. However, given the metamorphic grade of rocks from the locality (Flannery et al., 2018) and the possibility of contamination from contemporary surface‐dwelling microbes, results such as this should be viewed with caution. Nevertheless, cherts formed on early Mars would not have been subjected to such intense metamorphism and may retain biosignatures in contrast to similar rocks on Earth (see section 5). Silicification of organic remains involves the bonding of silicic acid to organic cell walls or envelopes, ensuring long‐term stability (e.g., Knoll, 1985). Experiments have confirmed that diverse archaeal and bacterial extremophiles and even viruses silicify readily in silica saturated solutions, with minimal dependence on cell/substrate type, pH, or salinity (Orange et al., 2009, 2013, 2014; Westall et al., 1995; Westall, 1997). Such results suggest that silicification could outpace cell lysis and degradative processes in brines on early Mars (e.g., Harrison et al., 2016; Toporski et al., 2002; Yee et al., 2003). Besides occurring in hydrothermal settings, amorphous silica is expected to be present on Mars as a result of low‐temperature chemical weathering of basalt (McLennan, 2003; Tosca et al., 2004), and orbital and in situ observations have shown it to be widespread (Milliken et al., 2008; Squyres et al., 2008; Sun & Milliken, 2015). It may be challenging to differentiate hydrothermal silica from silica enrichment by in situ weathering processes (potentially including “acid fog”; Tosca et al., 2004); such weathered rocks may be aggressively altered and are not expected to preserve biosignatures. Silica‐rich deposits near the remotely observed Nili Patera caldera have been suggested to represent ancient hydrothermal systems, although not necessarily formed at the surface, based on their setting and distribution (Skok et al., 2010). Siliceous material examined by the Spirit rover in Gusev Crater has also been interpreted as a hydrothermal deposit, although other explanations of its chemical composition are possible (Squyres et al., 2008). Stratiform, “rubbly” nodules of opaline silica in Gusev Crater have been considered morphologically comparable to digitate, biologically influenced sinter nodules produced in shallow water at the El Tatio volcanic spring in Chile, which are rich in preserved microbial filaments (Ruff & Farmer, 2016). However, the simple digitate appearance of these nodules could arise from abiotic processes (aggregation, concretionary growth, sedimentation, and/or weathering) and is not in itself a biosignature or even definitive evidence of hot spring deposition (Anderson, 1930; Grotzinger & Knoll, 1999; McLoughlin et al., 2008). Nevertheless, true hot spring sinters on Mars would represent an excellent search target for silicified microfossils and organic matter from a paleoenvironment likely to have been habitable. Perhaps, the most promising location for preservation in silica identified to date on Mars occurs at Gale Crater, where a 5–10‐m‐thick interval of lacustrine strata in the Murray Formation is enriched in silica (see below).

2.5 Chert and Silicified Sediments Bedded and nodular marine cherts on Earth, which are a diagenetic product of amorphous silica precipitated at or just below the seafloor, represent a major source of well‐preserved microfossils, particularly of Precambrian age (e.g., Barghoorn & Tyler, 1965; Schopf, 1968; Schopf et al., 2008) when pore waters became silica saturated in the absence of silica‐secreting organisms (Maliva et al., 2005; Siever, 1992). This led to rapid, early (perhaps syndepositional) silica precipitation that formed a rigid, impermeable solid material, resistant to later fluid alteration (Bartley et al., 2000; Ramseyer et al., 2013; Stolper et al., 2017). Such silica‐rich rocks on Earth, especially where amorphous content is high, can result in good or even spectacular preservation of cells and colonies (Figures 1e and 1f); dozens of examples are known from the Proterozoic (Schopf & Klein, 1992), including the iconic, densely packed assemblages of filamentous and coccoidal bacteria in the Gunflint Formation (Ontario, Canada, ~1.9 Ga) and the Bitter Springs Formation (Central Australia, ~850 Ma). The Archean record is more haphazardly preserved, probably because of ubiquitous hydrothermal and metamorphic overprinting and recrystallization. Chert (including early replacive chert after carbonate) is the dominant preserving medium of most purported microfossils older than 2.5 Ga (Schopf, 2006, and references therein), although most of these are controversial (Brasier et al., 2006). Historically, however, there has been a lack of attention to siliciclastic lithologies, which have recently begun to prove fruitful (Javaux et al., 2010; Wacey et al., 2011). On Mars, the aqueous alteration of basaltic crust is thought to have supplied abundant silica to rivers and lakes (McLennan, 2003; McLennan & Grotzinger, 2008). As on Earth, this silica could have solidified very early, providing an ideal medium for the preservation of any microorganisms living in the water column, on the lake floor, or in shallow subsurface sediment. Opaline silica has been observed from orbit in laterally continuous, well‐stratified deposits adjacent to the Valles Marineris canyon system, and in some cases these deposits occur as inverted channel systems (Milliken et al., 2008; Weitz, Milliken, et al., 2008; Weitz et al., 2010). Silica has also been observed in strata associated with potential sublacustrine fans within Melas Chasma (Metz et al., 2009) and in closed basins in the Noctis Labyrinthus region (Thollot et al., 2012). More recently, the Curiosity rover recovered evidence for sedimentary silicification in silica‐rich mudstones in the Marias Pass area of Gale Crater where the lower Murray Formation forms a thick sequence of lacustrine mudstones that interfingers and overlies fluvial‐deltaic sandstones and conglomerates (Grotzinger, Gupta, et al., 2015; Morris et al., 2016; Hurowitz et al., 2017; Rampe et al., 2017). The grain size is below the limit of resolution (<60–70 μm) of the Mars Hand Lens Imager, and parallel stratification with a mean lamina thickness of about ~0.5 mm extends laterally for at least several tens of centimeters (Figure 2). Quiescent subaqueous deposition is further evidenced by the absence of cross stratification, mudcracks, or any evidence for transport, erosion, or reworking. The “Buckskin” rock drilled by the Curiosity rover is characterized by ~40 wt % crystalline and ~60 wt % X‐ray diffraction amorphous material, and a bulk composition of ~74 wt % SiO 2 . The crystalline silica comprises trydimite and cristobalite with a bulk rhyolite‐like composition, suggesting a felsic volcanic provenance for the sediment (Morris et al., 2016). The amorphous material is silica rich, ~39 wt % opal‐A and/or silica glass and opal‐CT, and most likely represents an authigenic lacustrine precipitate or diagenetic alteration product (Hurowitz et al., 2017; Morris et al., 2016). Figure 2 Open in figure viewer PowerPoint Photograph taken by the Curiosity rover of “Lamoose” target, a float block from the Murray Formation, Gale Crater, Mars. It is minimally dust covered (reddish tone) and sculpted by wind to reveal very fine lamination (here oriented upper left to lower right) and fine grain size. Wind‐induced surface striations trend obliquely to the primary depositional lamination. Scale bar: 1 cm. Such a facies of finely laminated, fine‐grained silica‐rich mudstone, with substantial amorphous silica, may represent a favorable context for microfossil preservation. Early lithification would have sealed the rock (and any contained fossils) from later fluids that oxidized other parts of the Murray Formation (Hurowitz et al., 2017; Rampe et al., 2017). The presence of magnetite of probable authigenic origin in one drill hole (rather than hematite as in 14 others spread over ~150 m of section) signifies a lower degree of oxidation in either the primary or diagenetic environment, or both (Grotzinger, Crisp, et al., 2015; Hurowitz et al., 2017; Morris et al., 2016; Rampe et al., 2017; Vaniman et al., 2014). Indeed, magnetite as well as some of the hematite in Gale Crater could have formed by redox‐related primary precipitation, a process conducive to the preservation of cellular fossils (Fraeman et al., 2016; Hurowitz et al., 2017). Iron oxides can also adsorb silica and enhance silica precipitation, further strengthening the potential for preservation (Meister et al., 2014). The occurrence of a silica‐magnetite mudstone facies in part of the Murray formation demonstrates the potential for finding similar types of sedimentary rocks in lacustrine settings elsewhere on Mars, providing a strong candidate for sample return.

2.6 Siliciclastic Sediments Orbiter‐obtained geomorphological evidence for siliciclastic facies on Mars indicates alluvial fan, fluvio‐deltaic, sublacustrine fan, and aeolian deposits (e.g., Dromart et al., 2007; Malin & Edgett, 2003; Metz et al., 2009; Milliken et al., 2014; Moore & Howard, 2005). Rover observations confirmed the presence of proximal to distal fluvial, deltaic, lacustrine, and aeolian facies (Grotzinger et al., 2005; Grotzinger et al., 2014; Lewis et al., 2008; Williams et al., 2013). The organization of fluvial systems at Gale Crater shows facies transitions analogous to terrestrial “source‐to‐sink” networks leading to accumulation of lacustrine mudstones hundreds of meters thick (Grotzinger, Gupta, et al., 2015; Szabó et al., 2015; Fedo et al., 2017). Fluvial channel sediments represent high‐energy depositional environments where the chances of preserving fossils are poor, although allochthonous biosignatures may result from reworking, transport, and sometimes concentration, for example, in fossiliferous clasts, by fluvial processes. Distal deltaic, lacustrine, shoreline and subtidal deposits, on the other hand, could preserve a wide range of sedimentary microbialites (Figures 1d and 1g), morphological fossils (Figure 1h), and/or organic biosignatures (e.g., Ehlmann, Mustard, Fassett, et al., 2008; Summons et al., 2011; Grotzinger et al., 2014). Siliciclastic sediments are texturally and chemically diverse and preserve fossils in a variety of modes. Fine‐grained and clay‐rich siliciclastic lithologies are associated with some of the best preservation of microbes and soft‐bodied eukaryotes on Earth (e.g., Butterfield, 1990, 1995; Callow & Brasier, 2009; Farmer & Des Marais, 1999; Javaux & Knoll, 2017; Yuan et al., 2011), including those of Archean age (Javaux et al., 2010). The low permeability of fine‐grained sediments limits diffusion away from decaying remains once they are buried and favors the precipitation of authigenic minerals such as carbonate, pyrite, and phosphate. These minerals can replicate cells and tissues and/or cement the grains around them into concretions or high‐resolution molds, a process that has been studied experimentally (e.g., McCoy et al., 2015). The charged surface area of sedimentary clay minerals adsorbs and retains organic matter; the organic carbon content of marine mud and ancient shales correlates strongly with the total surface area of clay minerals within these lithologies (e.g., Hedges & Keil, 1995; Kennedy et al., 2002), particularly in sediments rich in smectite (Ransom et al., 1998). Cyanobacteria become coated with clay minerals in less than a week in experiments with sand, silt, dissolved silica, and suspended clays (Newman et al., 2016, 2017), ultimately resembling Precambrian‐Cambrian fossil filaments composed of aluminosilicates (e.g., Callow & Brasier, 2009). However, the relative importance of trapping suspended clays versus clay precipitation in natural environments remains unknown, and trapping is thought to dominate (Konhauser et al., 1998; Konhauser & Urrutia, 1999; Newman et al., 2016, 2017). The late Precambrian and early Paleozoic fossil record includes a large number of clay‐hosted Konservat‐Lagerstätten that preserve soft tissues as carbonaceous compressions, commonly with a secondary coating of authigenic clays that appears to track the original organic matter (Briggs, 2003). The role of preexisting clay in retarding decay appears to be more important for this style of preservation than the precipitation of early authigenic minerals (Gaines et al., 2008). Al3+ and Fe2+ ions, for example, may stabilize organic matter by promoting the crosslinking (“tanning”) of proteins or inhibiting the activity of autolytic enzymes (Butterfield, 1995; Petrovich, 2001; Wilson & Butterfield, 2014). Experiments have shown that polychaetes and crustaceans buried in the aluminum‐rich clay kaolinite for months‐to‐years are better preserved than those buried in other minerals (Wilson & Butterfield, 2014; Naimark, Kalinina, Shokurov, Boeva, et al., 2016). Clays rich in Al3+ and Fe2+ have likewise been shown to suppress the growth of various heterotrophic bacteria, including representatives of the microbial community typically involved in tissue decay, providing the first clear evidence of how clays might inhibit decay (McMahon et al., 2016; Morrison et al., 2016). Such interactions may explain why particular clay mineralogies correlate with the presence or absence of fossils in some stratigraphic sections (e.g., Anderson et al., 2014, 2018). More generally, taphonomic factors identified as favorable for preservation in siliciclastics include reduced microbial activity in sediments, the presence of iron, the type of clay, the activity of microbes in photosynthetic mats, elevated concentrations of silica, redox gradients, and the activity of sulfate reducing microbes that ultimately produce pyrite (Darroch et al., 2012; Gehling, 1999; Laflamme et al., 2011; Naimark, Kalinina, Shokurov, Boeva, et al., 2016; Tarhan et al., 2016; Wilson & Butterfield, 2014). The effects of most of these factors need to be investigated further before we can infer their likely impact on Mars. However, our current understanding suggests that Fe/Mg‐rich detrital smectites in fluvio‐deltaic and lacustrine deposits in ancient lake basins on Mars provide a promising context for fossil and organic preservation, especially where Fe‐rich end‐members can be identified (e.g., Ehlmann, Mustard, Fassett, et al., 2008; Hurowitz et al., 2017; Milliken & Bish, 2010; Rampe et al., 2017; Vaniman et al., 2014). Textural evidence of microbial activity occurs in siliciclastic lithologies from mudstone to sandstone and ranges from patterned textures on bedding planes (“microbially induced sedimentary structures” or “MISS”; Noffke et al., 2001) to three‐dimensional stromatolites (Schieber et al., 2007). In the absence of direct evidence of biogenicity, some MISS are difficult to distinguish visually from structures formed by sediment loading, shrinkage or shearing, or by the interaction of sediment with currents, escaping pore water, or early cements (Davies et al., 2016; Schieber et al., 2007). Others can probably only form via the growth of microbial mats on sandy surfaces, which may wrinkle (Figure 1g) or crack subaqueously at scales that indicate microbial aggregation or biostabilization (Gehling & Droser, 2009; Mariotti, Perron, et al., 2014; McMahon et al., 2017). Interactions between microbial surfaces, clay minerals, and microbial sulfide or silica at the surface of some MISS can preserve organic matter and replicate ~100‐μm‐scale filamentous microfossils in clay minerals or pyrite, even in relatively coarse lithologies (Callow & Brasier, 2009). MISS could be recognizable at distances of several meters and would warrant investigation if detected. However, the hypothesis that such structures may be visible in Curiosity images from the Gillespie Lake sandstone member within Gale Crater (Noffke, 2015) is not compelling. In common with Davies et al. (2018), we interpret the photographed features as erosional/fracture surfaces, not bedding planes; where dust‐free bedding planes are exposed on nearby outcrops, they show no textural features attributable to microbial mats. Indeed, MISS may be difficult to recognize on Mars because exposed bedding planes tend to be effaced by aeolian weathering.