Icarus 211 (2011) 1204–1214
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Icarus
journal homepage: www.elsevier.com/locate/icarus
Sequence and timing of conditions on early Mars
Caleb I. Fassett ⇑, James W. Head
Department of Geological Sciences, Brown University, 324 Brook Street, Box 1846, Providence, RI 02912, USA
a r t i c l e
i n f o
Article history:
Received 21 July 2010
Revised 5 November 2010
Accepted 9 November 2010
Available online 19 November 2010
Keywords:
Mars
Mars, Surface
Geological processes
Astrobiology
a b s t r a c t
The geological record of early Mars displays a variety of features that indicate fundamental differences
from more recent conditions. These include evidence for: (1) widespread aqueous alteration and phyllosilicate formation, (2) the existence of an active magnetic dynamo, (3) the erosion of extensive valley networks, some thousands of kilometers long, (4) a much more significant role of impact cratering, forming
structures up to the scale of large basins, and (5) the construction of much of the Tharsis volcanic province. Mars also is likely to have had a much thicker atmosphere during this early period. We discuss and
review the temporal relationships among these processes and conditions. Key observations from this
analysis suggest the following: (1) the last large impact basins, Argyre, Isidis, and Hellas, all pre-date
the end of valley network formation, potentially by several hundred million years, (2) the magnetic
dynamo is likely to be ancient (pre-Hellas), since the center of Hellas and other young basins lack magnetic remanence, and (3) the period of phyllosilicate formation is not readily connected to the period of
valley network formation. Concepts for the possible formation and evolution of life on Mars should
address this time sequence of conditions.
Ó 2010 Elsevier Inc. All rights reserved.
1. Introduction
Spacecraft data indicate that the early environment of Mars differs from recent conditions in a multitude of important ways (see,
e.g., Solomon et al., 2005; Carr and Head, 2010). Before the midHesperian, Mars appears to have had higher impact flux (Hartmann
and Neukum, 2001), a wetter surface (e.g., Carr, 1996; Craddock
and Howard, 2002), more volcanic resurfacing (Tanaka et al.,
1987), neutral-pH aqueous alteration (Bibring et al., 2006; Murchie
et al., 2009), an intense magnetic dynamo (Acuña et al., 1999), and
possibly a denser atmosphere (e.g., Jakosky and Philllips, 2001).
More speculatively, Mars may have had an ocean on its Noachian
surface (Baker et al., 1991; Clifford and Parker, 2001; Di Achille
and Hynek, 2010); direct evidence strongly favors the existence
of many large lakes (see, e.g., Irwin et al., 2002; Fassett and Head,
2008a). Each of these factors, with the possible exception of a higher impact flux, is broadly consistent with a more habitable Mars in
the Noachian to Early Hesperian than at present. The potential habitability of the ancient planet has helped motivate an exploration
strategy predicated on examining geological materials from this
early period (e.g., Grotzinger, 2009).
However, given the fact that conditions on early Mars appear
distinct from those observed today, it is common to assume that
there is a discrete geological period perhaps of some length when
all of these conditions existed simultaneously (active magnetic
⇑ Corresponding author.
E-mail address: [email protected] (C.I. Fassett).
0019-1035/$ - see front matter Ó 2010 Elsevier Inc. All rights reserved.
doi:10.1016/j.icarus.2010.11.014
field, valley formation, erosion and transport, aqueous alteration,
etc.). Although such a scenario is possible, a variety of observations
constraining the timing of these processes suggests that it may not
be the most probable scenario. In this paper we review the constraints on the timing of various conditions based on stratigraphy,
crater counting, inferences from the martian meteorite ALH84001,
and a variety of orbital observations.
1.1. Crater statistics and the age of surfaces and materials
Given our current lack of samples acquired from known locations on Mars, the primary technique for deriving ages of surfaces
or geomorphic features is to measure their superposed crater sizefrequency distribution (e.g., Hartmann, 1966; Soderblom et al.,
1974; Neukum and Wise, 1976; McGill, 1977; Tanaka, 1986;
Barlow, 1988, 1990; Strom et al., 1992; Hartmann and Neukum,
2001; Neukum et al., 2010). Since craters are presumed to accumulate in a spatially random process, at least insofar as the crater population is dominated by primary impactors, areas with higher
spatial densities of craters are interpreted to be older. If a reasonable model for the rate at which craters are accumulating can be
obtained (e.g., Hartmann and Neukum, 2001), absolute ages can
be estimated. These are model dependent, and are therefore less
definitive than relative age determinations.
A complicating factor in using craters to derive relative or absolute ages is that the number of craters observed in a given region is
not independent of its geological history. Gradation, erosion, and
exhumation can remove craters from the surface population or
C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
expose the remnants of craters from an earlier era (e.g., Grant and
Schultz, 1990; Malin and Edgett, 2000; Hartmann and Neukum,
2001). The fact that crater counting measurements are a combination of surface age and crater retention mediated by geological processes has been emphasized by Hartmann (1966), who proposed
the useful concept of ‘‘crater retention age.’’ Crater retention ages
of a given surface can vary substantially (by an order of magnitude
or more) when examining craters in different size ranges or at different reference diameters. The crater population at larger diameters is more representative of unit emplacement ages than the
population of smaller craters (<1–2 km), which are easier to
remove from the record and commonly reflect rates of regolith
processes rather than the age of underlying bedrock (e.g.,
Hartmann, 2005). If we are primarily interested in the age of
emplacement for surface units, this imposes practical limitations
on the minimum area where formation age can be ascertained, since
large craters form considerably less frequently than small craters.
However, since many of the major stratigraphic markers of important events in early Mars history subtend large areas, such as large
impact basins, this limitation does not preclude developing a reasonably strong relative sequence of events from crater statistics alone.
2. Basin and impact flux, and basin sequence
The impact flux at Mars and throughout the inner Solar System
is inferred to have been relatively constant over the last 3 Gyr, perhaps within a factor of 1.5 of the modern rate, but prior to that time
is thought to have increased rapidly (e.g., Hartmann, 1972;
Neukum and Wise, 1976; Guinness and Arvidson, 1977; Neukum
et al., 2001). At the Noachian–Hesperian boundary the impact rate
may be a factor of 80 greater than rates at present (Neukum et al.,
2001; Hartmann and Neukum, 2001). Absolute age estimates for
the Noachian/Hesperian boundary are TNH = 3.5–3.75 Gyr, depending on the model age system being used (see Hartmann and
Neukum, 2001; Ivanov, 2001; Hartmann, 2005; Fassett and Head,
2008b). Earlier, the rate of impacts is unknown and depends on
numerous factors, including whether there was a focused period
of heavy bombardment that affected Mars at 3.9 Gyr (see
Chapman et al., 2007).
The flux model of Hartmann and Neukum (2001) assumes that
there was no impact ‘spike’ per se, but impacts at 3.9 Gyr are still
presumed to be a factor of 3 greater than at 3.74 Gyr (and a factor
of 250 times the present rate). If the early impact flux was sufficiently high during the first 600 Myr of Mars history (>4 Gyr), it is
possible that the terrain may have been cratered to saturation
equilibrium, a condition where on average every new crater erases
a pre-existing crater of comparable size. Recent modeling and
observational evidence suggests that this condition was achieved
on the Moon (Richardson, 2009; Head et al., 2010). On Mars, terrain
from this period where the crust was saturated with impacts may
be described as ‘‘Pre-Noachian’’ (Frey et al., 2003; Nimmo and
Tanaka, 2005); the pre-Noachian/Noachian boundary is assigned
to the Hellas impact. Even when saturation is reached, preservation
of features on the surface is likely to be scale-dependent: the
signature of large impact basins (>500 km in size), such as their
long-wavelength topography, persist even when other geological
features from this time are entirely obliterated. An example of such
a persistent topographic signature is the martian dichotomy boundary, since it pre-dates the oldest mappable surface units and demarcates a major difference in crustal thickness (Solomon et al., 2005).
Much of the rock mass making up the martian crust probably also
pre-dates the surface age of materials that are exposed, as recorded
in the ancient age of the meteorite ALH84001, 4.09 ± 0.03 Gyr
(Lapen et al., 2010) (note that earlier estimates imply an age for
ALH84001 closer to crustal formation; 4.50 ± 0.13 Gyr; Nyquist
et al., 1995).
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Because of this much enhanced impact flux on the early planet,
large basin formation (impacts producing craters with diameter
>500–750 km) appears to have occurred only during the early history of Mars, from its formation until the end of the Noachian. This
observation may be slightly at odds with the models for crater flux
at Mars that are usually applied. Using Poisson statistics and the
cratering models of Hartmann (H; Hartmann, 2005) and Neukum
(N; from Ivanov (2001)), we calculate a large chance of a basin later
in Mars history: a 60% (H) and 72% (N) chance of at least one crater
750 km or larger since the Noachian/Hesperian boundary and an
89% (H) and 99% (N) chance for craters larger than 500 km. These
calculations would imply an expected number of craters
P500 km between 2 (H) and 4 (N) since the end of the Noachian.
Although small number statistics (N = 0) are unreliable, the lack
of any obvious candidates for post-Noachian impact basins larger
than 500 km suggest that the formation rate of very large craters
may be overestimated late in Mars history by these models. This
argument assumes that if the expected basins formed during the
post-Noachian, they would stand a good chance of surviving as recognizable impact structures, which is plausible since the rates of
crater modification and erosion in the Hesperian and Amazonian
on Mars appear low (e.g., Craddock and Maxwell, 1993; Golombek
et al., 2006).
This apparent absence of late large basins is in agreement with,
though does not prove, the hypothesis that the impactor size population changed at 3.8 Gyr (Strom et al., 2005), perhaps related to
the end of the Late Heavy Bombardment. Strom et al. (2005) argue
that this change in impactor population is reflected by an increased
frequency of larger impacts (greater than 10–20 km) compared to
smaller impacts (10–20 km and smaller) prior to the end of the
heavy bombardment on old terrains of the Moon, Mercury, and
Mars compared to younger plains. Recent observations consistent
with this change have been made on the Moon (Head et al.,
2010), Mercury (Fassett et al., manuscript in preparation) and
Mars, where Werner (2008) noted that it appears that ‘‘[t]he
basin-forming projectile population is most likely different from
the general impactor population’’.
Regardless of the uncertainties in the cratering flux and impactor populations, it is possible to use the superposed visible crater
populations to directly assess the formation of the largest wellexposed basins on Mars in relative terms. Independent crater
counts on the best preserved rim regions of Argyre, Isidis, Hellas
and other basins by Schultz and Rogers (1984), Werner (2008),
and us, all suggest that the sequence of the largest, well-preserved
impact basins was Hellas, Isidis, then Argyre (see Table 1 and
Fig. 1); note that Tanaka (1986) would have Isidis before Hellas.
Based on our data, Hellas is at the base of the Noachian, and Isidis
and Argyre are Early-to-Mid Noachian.
Along with their sequence, basins represent important stratigraphic markers on the surface which can be used to directly infer
environmental changes. One factor that is important is that all of
these basins have been incised by valley networks on their interior
or immediate exterior, implying that substantial fluvial activity
took place after their formation (Fig 1). This is consistent with crater counting evidence that valley networks continued to be active
until at least the end of the Late Noachian or possibly into the Early
Hesperian (Fassett and Head, 2008b). However, Argyre appears to
have the best preserved basin-related facies (e.g., Schultz, 1986;
see also Fig. 1); thus, the inferred sequence of basins from crater
statistics is also supported by the preservation state of the basins.
Several recent studies have extended the search for the visible
crater population to use topography and gravity data to map quasi-circular depressions (QCDs), circular crustal thickness anomalies
(CTAs), and ghost craters, which are interpreted as buried or highly
degraded impact structures superposed on many martian surfaces
(Frey et al., 2002; Head et al., 2002; Frey, 2006, 2008; Edgar and
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C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
Table 1
Crater measurements of the Hellas, Isidis, and Argyre. N(X) is the cumulative number of craters PX, normalized to an area of 106 km2; errors in N(X)
p
are from ±r = N/A. The basins are clearly distinguishable from each other on the basis of the relative density differences for craters larger than
20 km. Age estimates are model ages from Hartmann (2005) isochrons (AH) and Neukum isochrons (AN) (reproduced in Ivanov (2001)). The
computed ‘model ages’ are very insensitive to changes in crater frequency because of the high flux of impacts assumed in these absolute age models
early in Mars history (as can be seen by comparing the frequencies and derived ages for Hellas and Argyre). Statistical fit errors for given model ages
are ±0.01 Gyr. In reality, the age is not known nearly this well; uncertainty in age estimates is dominated by the systematic uncertainty in the
absolute age calibration/impactor flux (see, e.g., Hartmann and Neukum, 2001; Werner, 2008). Count areas are shown in Fig. 1 along with the crater
size-frequency distributions for Hartmann and Neukum isochrons.
Basin
Hellas
Isidis
Argyre
Count area (km2)
5
7.3 10
4.7 105
1.6 106
N(20)
N(64)
AH
AN
Period
151 ± 14
117 ± 16
88 ± 7
27 ± 6
11 ± 5
10 ± 2.5
4.02
3.96
3.92
4.04
3.97
3.94
Base of the Noachian
Mid-to-Early Noachian
Mid (to Early?) Noachian
Frey, 2008). The recognition of quasi-circular depressions and
ghost craters has been particularly important for inferring that beneath the upper few hundred meters of the northern plains, there
is an older (Hesperian?) ridged plains unit that is not deeply buried
(Withers and Neumann, 2001; Head et al., 2002) and that below
this surface, the basement of the northern plains are as old as
the Noachian-aged highlands (Frey et al., 2002).
More recently, Frey (2008) and Lillis et al. (2008a) suggested that
the combined population of possible QCDs and CTAs may provide
the best data for considering the age of the largest basins (using features of 300 km and more in diameter). However, we are cautious
about relying on this approach, since their data would imply that
Argyre is older than Isidis by an appreciable margin, and that Isidis
is as young as the beginning of the Late Noachian (3.8 Gyr).
This sequence of basins (Hellas, Argyre, Isidis) is in direct disagreement with the measured population of visible craters (Table 1
and Fig. 1). Across a wide range of visible crater diameters (Table 1
and Fig. 1), there are fewer superposed craters on Argyre than
Isidis, and fewer craters on Isidis than Hellas. Moreover, the rim
region of Isidis (the Libya Montes) has a crater population dating
to the Mid-to-Early Noachian boundary, which is inconsistent with
a Late Noachian formation for the basin. Finally, the inferred
sequence of these youngest fresh impacts is inconsistent with the
relative preservation state of the basins (Schultz, 1986). We suggest
that possible reasons for this discrepancy include one or more of the
following factors: (1) small number statistics: since these young and
well-preserved basins never had many >300 km craters to begin
with (visible, degraded, QCDs or CTAs), inferring their relative age
from craters of this size may lead to errors, (2) some QCDs or CTAs
may not be impact structures, and/or (3) there may be different degrees of basin floor filling that affects the number of QCDs/CTAs that
can be recognized (Isidis is potentially more filled than Argyre)
(Head et al., 2002; Howenstine and Kiefer, 2005).
In summary, for the youngest, well-exposed basins on Mars, we
prefer to rely on the superposed visible crater population for
assessment of their timing and relative sequence – first Hellas,
then Isidis, then Argyre.
3. Valley networks and surface erosion
Valley networks provide morphological evidence for fluvial
activity, erosion, sedimentary transport, and a hydrological cycle
on early Mars (Carr, 1996). Valley networks have numerous tributaries (Hynek et al., 2010), often begin near drainage divides
(Craddock and Howard, 2002), and were interconnected across
great distances, at least during their period of peak activity (see,
e.g., Irwin et al., 2005; Fassett and Head, 2008a). Paleo-lakes on
Mars appear to have been relatively common features (e.g., Fassett
and Head, 2008a, and references therein), and certain valleys such
as Ma’adim Valles, which initially appeared to come from localized
sources (e.g., Gulick, 2001), appear to have formed as these
paleo-lakes overtopped confining topography (Irwin et al., 2002).
Groundwater-driven valley erosion alone seems inconsistent with
many valley characteristics, particularly the dendritic, high-order
tributaries that extend to drainage divides (Hynek et al., 2010). Even
if some valleys formed as the result of groundwater discharge, precipitation-based recharge seems to have been necessary to close the
hydrological cycle, as basic calculations suggest that subsurface
water reservoirs would need to be recharged many times to erode
the valley networks observed (Goldspiel and Squyres, 1991; Gulick,
2001). The characteristics of valley networks thus seem to require,
at minimum, time periods when precipitation on the surface was
possible, water was cycled through the early Mars atmosphere,
and water was stable or metastable at the martian surface
(Craddock and Howard, 2002; Hynek et al., 2010).
Several independent studies have attempted to estimate when
the most extensive period of valley network formation occurred
(e.g., Pieri, 1980; Carr and Clow, 1981; Fassett and Head, 2008b;
Hoke and Hynek, 2009), using stratigraphic and crater counting
analysis to date the termination of valley network activity. These
studies suggest that regional-to-global-scale valley formation persisted until approximately the Noachian/Hesperian boundary or
into the Early Hesperian at the latest (Fassett and Head, 2008b;
Hoke and Hynek, 2009). Note that this ‘regional-to-global’ scale
formation excludes certain regions that are thought to be local
exceptions, such as valleys on certain volcanoes (e.g., Gulick and
Baker, 1990; Fassett and Head, 2006, 2007), in association with glaciation (Dickson et al., 2009; Fassett et al., 2010), and within, or in
the vicinity of, young, large craters (e.g., Williams and Malin, 2008;
Tornabene et al., 2008; Morgan and Head, 2009).
In our study (Fassett and Head, 2008b), we suggested that two
possible interpretations were consistent with our craters statistics:
either (1) global termination of valley activity near the Noachian/
Hesperian boundary or (2) persistence of some valleys into the
Early Hesperian.
Increasing evidence has been put forth for erosion in at least
some major valley networks were active well into the Early
Hesperian or possibly the Late Hesperian (e.g., Mangold and Ansan,
2006; Ansan and Mangold, 2006; Bouley et al., 2009, 2010). In
some of this work, a younger period of activity is derived than in
Fassett and Head (2008b), primarily due to differences in analytical
choices, particularly: (1) how count regions are aggregated, (2) different stratigraphic interpretations and, most importantly, (3) the
effective diameter used to compare observed crater populations
with isochrons. At some level, these factors are coupled, since larger diameter craters require greater aggregation of area to achieve
meaningful statistics, at the expense of the ability to discern real
local variation if it exists (as noted by Bouley et al. (2010)). As described above, reliance on smaller craters may result in younger
ages due to crater retention. In summary, age data continue to support the idea that regional to global-scale valley network formation
terminated in the Early Hesperian, although new evidence has
bolstered the interpretation that valley formation lasted into this
period (Bouley et al., 2009, 2010).
C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
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Fig. 1. (Top two rows) Crater size-frequency measurements in incremental (top row) and cumulative (second row) plots on terrains related to Argyre, Isidis, and Hellas (third
row), which result in Early-to-Mid-Noachian ages. These basins have 3–5 the crater density superposed on valley networks, which in aggregate have a frequency near the
Noachian/Hesperian boundary, valley network formation mostly terminated in the Early Hesperian. The bottom row shows examples of valleys superposed on each of these
major, young impact basins.
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C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
Given the observations on the timing of valley network formation, three key questions remain to be definitively addressed: (1)
how active was the period of valley network formation in the Noachian to the Early Hesperian, (2) how common/continuous were
the periods when valleys were forming on early Mars, and (3)
was water required to be stable over an extended period of time?
Estimates from modeling of valley network-associated sedimentary deposits imply emplacement times that are geologically
quite short, of order 1–1000 years (Jerolmack et al., 2004;
Kleinhans, 2005; Lewis and Aharonson, 2006; Kraal et al., 2008;
Kleinhans et al., 2010). However, these estimates are based on continuous activity and sediment transport; estimates that assume
terrestrial-like intermittency or sediment supply unsurprisingly
infer much longer periods of time (Moore et al., 2003; Fassett
and Head, 2005).
Drainage basin characteristics provide some of the strongest
arguments for valley network formation over an extended period
of time (>105 years) (Barnhart et al., 2009). Barnhart et al.
(2009) synthetically reconstructed pre-erosion topography of the
Parana drainage basin, and applied a variety of erosion scenarios
to examine their consistency with the topography we actually
observe. They found that an intense period of fluvial erosion and
precipitation lasting 103–104 Earth years would be sufficient to
erode the valleys that are observed. However, these intense erosion
scenarios resulted in a pattern of erosion far more integrated (with
more crater rims breached) than what we observe on the surface.
Thus, models that have greater episodicity, with runoff distributed
over 105–106 Earth years, are interpreted to be more consistent
with the drainage pattern observed in the Parana region (Barnhart
et al., 2009).
4. Volcanism
Volcanism is known to be a major factor in the long-term climate evolution of Mars, as eruptions liberate volatile species from
the planetary interior to the atmosphere (e.g., Jakosky and Philllips,
2001; Phillips et al., 2001; Craddock and Greeley, 2009). For this
reason, volcanism has commonly been inferred to be closely linked
to changes in the surface environment. The formation of the mass
of Tharsis in particular has been implicated in a transition from a
phyllosilicate-forming era (phyllosian) to a sulfate-forming era
(theiikian) (Bibring et al., 2006; Bibring and Langevin, 2008).
Constraining the timing of Tharsis volcanism is critical to
understanding whether this conclusion is reasonable. On the basis
of the fact that its emplacement and load on the Mars lithosphere
influenced the orientation of Late Noachian/Early Hesperian valley
networks, Phillips et al. (2001) argued that the bulk of the Tharsis
volcanic was emplaced during the Noachian. Further evidence that
Tharsis construction is ancient also comes from mapping, crater
counts, and analysis of the tectonic record (Plescia and Saunders,
1982; Anderson et al., 2001), as well as from observations that portions of Tharsis are magnetized, even at high elevations (7 km)
(Johnson and Phillips, 2005).
On the other hand, the interpretation that the bulk of Tharsis is
Noachian has been disputed by Craddock and Greeley (2009), who
point out that the lack of craters on much of Tharsis means that
most of its surface is Hesperian or Amazonian, and requires significant post-Noachian resurfacing. Craddock and Greeley (2009) estimate that lava deposits up to 10 km in thickness are required to
remove a sufficient number of craters to reset the terrain age.
It is plausible that these two views can be reconciled in a scenario where the majority of the crust at Tharsis is constructed in
the Noachian (crustal thickness 50–100 km; Neumann et al.,
2004), but where extensive volcanic resurfacing persists through
Hesperian and Amazonian times (see also Solomon and Head,
1982). However, the observation that a substantial amount of
Tharsis-building is ancient (e.g., back to the Mid-Noachian or
before) remains credible, as the existence of ancient, Noachian
regions is clear, particularly in the Thaumasia highlands (Plescia
and Saunders, 1982). Given that the magnetization of parts of Tharsis (Johnson and Phillips, 2005), early volcanism in these regions
may pre-date Hellas (see Section 6). The interpretation that the
construction of Tharsis near the end of the Noachian led to secular
changes which caused Mars to transition from a planet where
phyllosilicate formation was common to one dominated by sulfate
formation (Bibring et al., 2006; Bibring and Langevin, 2008) may
not be consistent with the fact the bulk of Tharsis may be old.
Hesperian and younger volcanism on Mars is also important
regardless of the timing of Tharsis. In particular, volcanic plains
emplacement, particularly focused in the northern lowlands, resurfaced 30% of the surface of Mars in this period (Head et al., 2002).
Estimates from Viking mapping suggests that more than half of the
volcanic resurfacing on Mars is Early Hesperian or younger (Tanaka
et al., 1987; Greeley and Schneid, 1991); higher resolution observations with recent data would imply that this is conservative, because small patches of volcanic plains have been increasingly
recognized in the highlands (Fassett and Head, 2008a).
In summary, the volcanic history of Mars should be closely correlated with a number of other conditions on the planet, including
the density of the atmosphere, atmospheric chemistry and volatile
inventory. As far as it can be determined however, the timing of
volcanism (e.g., Tanaka et al., 1987) does not imply a one-to-one
link between volcanism and surface conditions. No evidence exists
that a declining volcanic fluxes correlates well with atmospheric
loss, or that periods of Noachian volcanism helped facilitate transient clement conditions. Instead, the Hesperian volcanic deposits
that resurfaced 30% of Mars are volumetrically significant and
strikingly uneroded. Based on our current understanding of the
timing of volcanic deposits, secular changes in volcanism or major
volcanic events can not be directly connected to transitions in surface conditions.
5. Aqueous alteration
The recognition of alteration products on Mars has been revolutionized by observations in the last decade across the electromagnetic spectrum, first in the thermal infrared by TES and THEMIS
(e.g., Christensen et al., 2001; Wyatt and McSween, 2002; Osterloo
et al., 2008), and more recently, in the visible to near-infrared, by
OMEGA (Gendrin et al., 2005; Poulet et al., 2005; Bibring et al.,
2006; Bibring and Langevin, 2008) and CRISM (Milliken et al.,
2008; Mustard et al., 2008; Ehlmann et al., 2009; Wray et al.,
2009).
These data have resulted in the recognition of at least ten distinct environments where aqueous alteration products are observed (Murchie et al., 2009). Based on observations of hydrated
minerals, particularly Fe–Mg phyllosilicates, it appears that neutral-pH alteration on Mars was an important process in the Noachian (Bibring et al., 2006). Murchie et al. (2009) examined the
stratigraphic constraints on these deposits; we independently have
reexamined these environments from a crater counting and stratigraphic perspective (Table 2). The time-stratigraphy of mineral formation in many of these environments is complicated. One of the
major issues is that in some of the outcrops where phyllosilicates
are observed, they are likely to be detrital (e.g., Ehlmann et al.,
2008a; Murchie et al., 2009; Milliken and Bish, 2010). The timing
of the aqueous alteration that resulted in the formation of these
clays is thus not preserved – their present state could reflect Early
Noachian formation and Late Noachian physical weathering, transport, and deposition.
Where minerals remain in situ (authigenic alteration), it is
easier to make inferences about the timing of the geochemical
C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
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Table 2
Time constraints on major aqueous alteration environments on Mars. The type outcrops that show aqueous alteration on Mars are modified after Murchie et al. (2009) (his
Table 3) and re-ordered in approximate chronological order based on our independent evaluation of the stratigraphic constraints on these outcrops from crater counts and local
relationships. Most neutral-pH alteration is conceivably quite old (before the Late Noachian), which may pre-date valley network activity, or at least the termination of VN
formation. Evaporite deposits/chemical precipitates may be more common late in Mars history (Bibring et al., 2006; Murchie et al., 2009).
Aqueous environments
Type area (s)
Timing of aqueous alteration for the type area (s)
Layered phyllosilicates
Nili Fossae
Pre-Noachian or Early Noachian. Most of the deep crustal alteration/
phyllosilicate is interpreted to be pre-Isidis (Mustard et al., 2007, 2009;
Mangold et al., 2007). Exhumation is important
Mawrth Valles
Pre- or Early-to-Mid Noachian. Age constraint on phyllosilicate bedrock
is from craters in the Mawrth Valles region that appear to post-date
the phyllosilicate-bearing material (Michalski and Noe Dobrea, 2007).
Exhumation is important (craters turned into knobs; phyllosilicates
are exposed from underneath an eroded caprock)
Deep phyllosilicates
Exposures by craters
in the highlands
Pre-Noachian or Noachian, Difficult to Constrain More Specifically.
Multitude of exposures in central peaks, rims, and walls are excavated
crust, so limits on timing are hard to come by. Formation/alteration
was conceivably at depth (e.g., Parmentier et al., 2008)
Carbonate-bearing outcrops
Serpentine-bearing outcrops
Nili Fossae
Early-to-Mid Noachian. Outcrops are associated with olivine units
(Ehlmann et al., 2008b, 2010) that are interpreted to be directly related
to the Isidis basin-forming event (e.g., Mustard et al., 2009)
Intracrater clay-sulfates
Columbus Crater
Mid-Noachian? (Late-to-Early). Rim of Columbus crater has
N(5) 440 ± 179 (N = 6), implying a (uncertain) Mid-Noachian age for
the interbedded clay and sulfates described by Wray et al. (2009)
Phyllosilicates in intracrater fans
Jezero Crater, Holden Crater,
Eberswalde Crater
Unconstrained. Presumably detrital. In the case of Jezero crater (e.g.,
Ehlmann et al., 2008a), the source of sediments includes Early (or Pre?)
Noachian phyllosilicates and Mid-to-Early Noachian carbonates in the
watershed
Plains sediments (chlorides)
Terra Sirenum
Late Noachian/Early Hesperian. Crater counting of the type area
suggests has a LN/EH-boundary age for the THEMIS ‘glowing’ terrain
(Osterloo et al., 2008). Chlorides are presumably evaporitic in origin;
associated phyllosilicates may be detrital
Meridiani-type layered deposits
Valles-type layered deposits
Meridiani Planum
Valles Marineris ILDs
Late Noachian to Hesperian. These sulfate-rich deposits retain craters
rather poorly. In Meridiani, sulfate plains clearly embay highlands and
have an Early Hesperian crater density, which is thus a minimum age
for the observed water–rock interaction. ILDs are likely Hesperian in
age based on stratigraphy and crater counting (Quantin et al., 2010)
Siliceous layered deposits
Plains above VM
Hesperian to Amazonian. Deposits are superposed on Late Hesperian to
Early Amazonian surfaces
Polar gypsum deposits
Basal unit and surrounding dunes
Unconstrained. Sand in dunes and basal unit; period of alteration is
unbounded
environments where these formed. From Murchie et al.’s (2009)
classification of distinct aqueous environments, the most likely
examples of outcrops with in situ aqueous mineral formation are
(1) deep phyllosilicates (common highlands exposures usually in
crater rims, central peaks, or ejecta; Mustard et al., 2008); (2) layered phyllosilicates (such as Mawrth Vallis; e.g., Poulet et al.,
2005); (3) certain carbonate-bearing outcrops (and, more recently
discovered, serpentine-bearing outcrops; Ehlmann et al., 2010) situated with their ultramafic precursors (Ehlmann et al., 2008b),
and, (4) environments with chemical precipitates or evaporites
(chloride-bearing plains sediments, Osterloo et al., 2008; hydrated
silica deposits, e.g., Milliken et al., 2008; layered sulfates such as
those found in Meridiani Planum and Valles Marineris; e.g., Gendrin et al., 2005; sulfates interbedded with phyllosilicates on a crater interior; e.g., Wray et al., 2009). For this final class in particular,
chemical sedimentation may be a result of groundwater-driven
interactions with the upper crust, rather than surface precipitation,
runoff, and weathering; preservation of jarosite at these locations
also suggests that long-term arid conditions existed after the
emplacement of these chemical sediments (Elwood Madden
et al., 2004, 2009).
Despite the fact that both valley networks and phyllosilicate
clays are predominantly in Noachian terrains, evidence that demonstrates that valley networks and these alteration products are
characteristics of the same environment and formed at the same
time is limited. Water–rock interactions that formed clays may
have mostly ended by the time of the Isidis impact in Nili Fossae
(Mustard et al., 2007; Mangold et al., 2007), and much of the observed neutral-pH alteration may have occurred in very ancient
times (Poulet et al., 2005; see also Table 2). If this is the case, the
phyllosilicates may be older than the Late Noachian to Early Hesperian valley systems where clay-bearing sediments were transported and deposited, such as in Eberswalde, Holden, and Jezero
craters (e.g., Ehlmann et al., 2008a; Milliken and Bish, 2010).
Along with broad global trends, there are also differences in the
character of aqueous alteration around the youngest large impact
basins Isidis (Mustard et al., 2007, 2009; Mangold et al., 2007;
Ehlmann et al., 2009) and Argyre (Buczkowski et al., 2010).
Buczkowski et al. (2010) observe that although iron/magnesiumbearing phyllosilicates are exposed within and by the Argyre basin
structure, less mineralogical diversity is present than in a comparable setting at Isidis. Buczkowski et al. (2010) interpret the alteration
minerals of Argyre as primarily pre-dating the basin-forming event,
which acted to expose pre-existing alteration products in the
Noachian crust. The greater diversity of alteration products in the
Nili Fossae area associated with Isidis requires multiple alteration
events in distinct weathering environments (Ehlmann et al., 2009).
This distinction is consistent with Argyre being younger than Isidis
(Section 2) and with a hypothesized global decline in neutral-pH,
high-water–rock ratio aqueous alteration as a function of time.
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In summary, a transition in the character of aqueous alteration
from widespread neutral-pH aqueous alteration to more localized
acidic aqueous alteration is suggested by observations of hydrated
minerals on Mars (Bibring et al., 2006; Bibring and Langevin, 2008;
Murchie et al., 2009). This paradigm of high-water–rock ratio alteration followed by more water-limited alteration later in Mars history (Bibring et al., 2006; Hurowitz and McLennan, 2007) seems
borne out by new data even as many new alteration minerals have
been recognized on the surface of Mars.
6. Magnetic anomalies and cessation of the magnetic field
Observations from the Mars Global Surveyor magnetometer
experiment demonstrated that there are crustal magnetic anomalies observed over much of the surface, with the strongest anomalies concentrated in the southern highlands (Acuña et al., 1999).
These crustal anomalies imply the existence of a core dynamo on
Mars early in its history. The existence of this magnetic field may
have played an important role in arresting the loss of the early
Mars atmosphere by solar wind sputtering (e.g., Jakosky and Philllips, 2001), as well as shielding the surface from energetic cosmic
rays (e.g. Molina-Cuberos et al., 2001).
There are two compelling constraints on the timing of the Mars
magnetic field. First, crustal magnetic anomalies are largely absent
in the interiors of Hellas, Argyre, Isidis, and Utopia, as well across
most of Tharsis and most volcanic edifices on Mars, with the exceptions of Hadriaca Patera (Lillis et al., 2006) and Apollonaris Patera
(Hood et al., 2010). The simplest explanation for the lack of magnetization in these basins and volcanoes is that they post-date the
cessation of the magnetic field. If this interpretation is correct,
the core dynamo must have ended in the pre-Noachian before
the formation of Hellas (Lillis et al., 2008a,b; see Schubert et al.
(2000) and Hood et al. (2010) for alternative interpretations of
the timing of the magnetic field). It has been suggested that the
formation of large earlier basins that are now buried, such as Utopia, may have contributed to this termination (Roberts et al., 2009).
Second, a further possible constraint on the timing of the magnetic field comes from ALH84001, which has remanent magnetization consistent with acquisition in a magnetic field caused by a
core dynamo with strength 0.1–10 the present Earth dynamo
(Kirschvink et al., 1997; Weiss et al., 2002). This interpretation is
preliminary, however, as it is not entirely clear whether the
magnetization in this meteorite was acquired from a dynamo or
from pre-existing crustal fields (Gattacceca and Rochette, 2004).
If it was from a core dynamo, and the age of ALH84001 is
4.091 ± 0.03 Gyr as recently suggested (Lapen et al., 2010), this
would provide direct evidence of the persistence of a magnetic
dynamo until 4.09 Gyr.
If these suppositions are correct, and ALH84001 preserves a
core field and Hellas formed after the dynamo ended on Mars, this
also bounds the formation of Hellas to after 4.09 Gyr (consistent
with crater counting model ages, Table 1). Regardless of the evidence from ALH84001, the lack of magnetization within Hellas
strongly suggests termination of the magnetic field before the
basin formed, well before the end of valley network formation.
The termination of the magnetic field before the valley network
activity in the Late Noachian/Early Hesperian is consistent with:
(1) crater counting results, which are imprecise but suggest a
potentially long gap between Hellas and the end of valley formation and (2) stratigraphy, which irrefutably demonstrates that valley formation continued after Hellas, but provides no information
about the length of time between Hellas and the end of valley formation. Thus, if a magnetic dynamo was playing an important
shielding role for the surface and/or atmosphere, the shield may
have been removed well before water stopped playing an impor-
tant geomorphic role on the martian surface (in contrast to the
timeline in Jakosky and Philllips (2001)).
One observation that complicates this scenario is the apparent
complex magnetization that is observed in other Mars meteorites
(e.g., Collinson, 1986; Collinson et al., 1997). Because the shergottites (180 Ma) and nakhlites (1.3 Ga) are much younger than
Hellas (e.g., McSween, 1994), this requires that when magnetization is observed in these younger meteorite samples, it was not
acquired by cooling in the presence of a dynamo. Other processes
that are plausible include shock magnetization (Cisowski and
Fuller, 1978), acquisition from the Mars crustal field, or by contamination by terrestrial fields. The alternative is that the interpretation that Hellas and other non-magnetized basins formed in the
absence of a core dynamo is wrong. Given our understanding of
the spatial distribution of magnetic remanence on Mars, the
post-dynamo acquisition of magnetization in these samples is
the simplest explanation, consistent with a scenario where the
‘‘SNCs [were] more likely magnetized during or after impact than
during the initial magmatic cooling’’ (Rochette et al., 2005). Recent
measurements of the nakhilite Yamato 000593 support this interpretation, consistent with the absence of a global magnetic field on
Mars when Yamato 000593 formed, 1.8 Gyr (Funaki et al., 2009).
Complicating the interpretation of the magnetic record further
is the fact that the observed pattern of crustal magnetization is
heterogeneous, with virtually all of the strong remanent crustal
magnetism observed in the southern hemisphere and with only
weak magnetic signatures north of the dichotomy boundary.
One explanation for this heterogeneity is that hydrothermal alteration may have been critical in establishing where magnetization
in the crust is observed today (Solomon et al., 2005). If hydrothermal alteration of the crust was preferentially concentrated in
low-lying regions, such as the largest impact basins and northern
lowlands, the lack of magnetic signatures in the large, young
impact basins may be a result of this demagnization process, even
if the active dynamo persisted after their formation (Solomon
et al., 2005).
Alternatively, the hemispheric difference in observed crustal
remanence may reflect a single-hemisphere dynamo (Stanley
et al., 2008), perhaps resulting from degree-one convection (e.g.,
Zhong and Zuber, 2001). A hemispheric dynamo does not affect
the overall constraints on timing, since Hellas and Argyre are
surrounded by crust with strong remanent magnetization, so the
single hemispheric dynamo should still have affected these basins.
Thus, in the absence of other modifying influences, the lack of magnetization in these basins would still imply that they post-date the
cessation of the magnetic field, even if the remanent magnetization
was a result of a one-hemisphere dynamo.
Lower-altitude measurements of the Mars crustal magnetic
field would be very useful to help test which scenario is the best
explanation for the observed magnetic anomalies (Langlais and
Amit, 2008).
7. Atmosphere and possible atmospheric loss
Direct constraints on both the density of the early Mars atmosphere and its loss are somewhat limited. Some invocations of
higher atmospheric pressure early in Mars history have been based
simply on the need to explain valley network formation (e.g.,
Pollack et al., 1987). Many such modeling efforts assume that surface conditions when valley networks were formed must have
been above 273 K (averaged over a Mars year), and investigators
have built various models with different atmospheric pressures
and constituents to explore how such a requirement might be
met (see, e.g., Haberle, 1998 and references therein).
Isotopic measurements provide the strongest indication that
the early atmosphere was substantially denser than today, perhaps
C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
by a factor of 10 or more, and subsequently removed (summarized
in Jakosky and Philllips (2001)). These isotopic measurements constrain the early atmosphere by comparing the abundance of lighter
isotopic species, which are more efficiently stripped away by sputtering and hydrodynamic escape, to heavier species. The factor of
>10 higher density atmosphere early in Mars history that these
observations require is a minimum estimate, since impact erosion
of the atmosphere (e.g., Melosh and Vickery, 1989) does not fractionate lighter and heavier isotopes and may also have been an
important factor in atmospheric loss. As described above, the end
of valley network formation clearly post-dates the period when
the largest basins were formed and the period of highest impact
flux; thus, the majority of valleys on the surface post-date the period when this mechanism would have been most effective.
Although the change in climate associated with the loss of a significant atmosphere may be correlated with the end of the period
of valley network formation on Mars (e.g., Pollack et al., 1987;
Jakosky and Philllips, 2001), this should be seen as an assumption
rather than an observation. On the basis of atmospheric argon
observations, Craddock and Greeley (2009) suggest that the martian atmosphere may have been similar to the present atmosphere
as far back as the Mid-Noachian. Moreover, atmospheric escape
rates due to solar wind interactions have been measured to be currently quite low, which suggests that even over billions of years
only a few millibars of CO2 and minimal water would be removed
from Mars atmosphere (Barabash et al., 2007). This very slow loss
is in agreement with theoretical expectations from atmospheric
photochemistry (Hodges, 2002).
On the other hand, models incorporating both impact erosion
and sputtering suggest the loss of 95–99% of the atmosphere
(Brain and Jakosky, 1998) since the beginning of the geological
1211
record (Early Noachian). Reconciling scenarios for atmospheric
state, climate, and surface erosion remains an important goal
for further research. Observations of the Mars atmosphere by
Mars Science Laboratory (e.g., Mahaffy et al., 2009) and later by
the Mars Atmosphere and Volatile Evolution Mission (Jakosky,
2008) should help address these questions during the next decade
by providing improved measurements of atmospheric isotope
ratios, trace gases, and interactions of the atmosphere with the
solar wind.
8. Synthesis
The observations of individual processes outlined above allow
us to draw some inferences on the most likely sequence for various
conditions (Fig. 2). We summarize the relationships between these
conditions here:
1. It is likely that the magnetic field responsible for the crustal
magnetism observed pervasively in the Noachian highlands
(Acuña et al., 1999) was: (a) still active at 4.09 ± 0.03 Gyr
(because of magnetization in ALH84001; Kirschvink et al.,
1997; Weiss et al., 2002; age from Lapen et al. (2010)) and (b)
terminated by the time of the formation of Hellas (Lillis et al.,
2008a,b). On the basis of our discussion of the magnetic field,
this is what we call the baseline scenario in Fig. 2.
Two other scenarios for the dynamo history are reasonable: if
the magnetization in ALH84001 post-dates a core dynamo (perhaps because it was acquired from pre-existing local crustal
remanent fields), or if ALH84001 is much older than recently
determined, an ‘‘early scenario’’ is possible. Or, if the lack of a
magnetic signature in Utopia, Hellas, Isidis and Argyre is not a
Fig. 2. A schematic of the sequence of various planetary conditions on Early Mars based on the information described in the text. Note that in this diagram we accept the
existence of a pre-Noachian period defined as the time before the Hellas impact, from which no known surface units date on the modern surface (Frey et al., 2003; Nimmo and
Tanaka, 2005). Along with the conditions we show, other important environmental conditions were (1) the general impact flux, which is thought to have declined since the
Late Noachian; before that time it may have had a peak (during the Late Heavy Bombardment) or simply a monotonic rise (see discussion in text), and (2) the atmospheric
density, for which the time-history is poorly understood, though evidence suggests that the atmosphere was denser during early periods than it is at present. The scenarios
for the timing and history of the core dynamo are particularly complex (see the text for more discussion and references). The baseline scenario shown here assumes that the
magnetization of ALH 84001 was frozen in an active dynamo; the ‘‘Early Scenario’’ would require its magnetization from pre-existing crustal fields. The ‘‘Late Scenario’’
discards the idea that the large basins post-date the magnetic field and requires a different explanation (e.g., thin crust, lack of magnetic carriers, hydrothermal alteration) for
why they lack apparent crustal magnetization. Evidence suggests that the beginning of the construction of Tharsis pre-dates the termination of the dynamo (Johnson and
Phillips, 2005) and that the bulk of the Tharsis load was in place by the period of valley network formation (Phillips et al., 2001).
1212
C.I. Fassett, J.W. Head / Icarus 211 (2011) 1204–1214
result of their formation after the core dynamo terminated (e.g.,
Schubert et al., 2000; Solomon et al., 2005; Hood et al., 2010), a
‘‘late scenario’’ is possible. In this case, the timing of the core
dynamo is not bounded except by the lack of a global magnetic
field from a dynamo today.
2. Hellas is plausibly younger than 4.09 ± 0.03 Gyr on the basis of
its non-magnetization and ALH84001’s magnetic signature. On
its face, this is consistent with crater model ages of 4.02–
4.04 Gyr (though the systematic calibration is far more uncertain than this range suggests). This upper bound on the absolute
age of Hellas goes away if the core dynamo terminated before
ALH84001 (the ‘early scenario’) or if the new ALH84001 age
(Lapen et al., 2010) is too young and it is actually older (as originally thought). Unless one of the hypothesized reasons requiring a ‘late scenario’ is correct, the relative age sequence where
the magnetic dynamo terminated before Hellas remains.
3. Large basins like Hellas, Isidis, and Argyre pre-date the end of
valley network formation, and hence the magnetic field likely
does so as well. The gap between the termination of the magnetic field and the formation of the Late Noachian to Early Hesperian valleys could be appreciable, depending on the absolute
length of the Noachian; current impact models would suggest a
period of 0.3–0.5 Gyr between the Hellas impact and end of
widespread valley formation. Neither the loss of the magnetic
field, nor a decline in the rate of volcanism or the impact rate,
connects in a one-to-one manner with the decline in valley network formation.
4. Similarly, formation of much of the phyllosilicate record that
indicates that pervasive aqueous alteration on Mars is difficult
to connect temporally to the period of valley network formation; many of the alteration products that are observed are
likely to be older than at least the last period of widespread valley formation.
5. Portions of Tharsis are magnetized (Johnson and Phillips, 2005),
suggesting that Tharsis construction began in the Early-to-Mid
Noachian or before. This is consistent with ancient tectonic
activity in parts of Tharsis (e.g., Plescia and Saunders, 1982)
and with the observation that the bulk of the Tharsis load was
in place before valley network formation (Phillips et al., 2001).
A secular change in the Mars environment linked to Tharsis formation cannot be connected in a one-to-one manner with
observations of the shift in the nature of aqueous alteration
environment.
A few other implications from these timing constraints are
apparent. As has been discussed before (Fassett and Head, 2008b;
Hynek et al., 2010), the obvious large basins (>500–600 km) on
Mars appear too old to be the direct cause of valley formation, in
contrast to the original scenario described by Segura et al.
(2002), where >100 km impactors lead to surface warming and valley formation. If the impact hypothesis described by Segura et al.
(2002) is to work, smaller impactors are more likely to be the cause
of valley networks (see also Toon et al., 2010). Timing constraints
alone allow for this possibility, although whether it is possible to
reconcile the observed erosion with the erosion that impacts might
produce still seems uncertain.
Second, these results suggest that if the magnetic field of Mars
was necessary for protecting life at surface of Mars, valley sediments and even phyllosilicates that date to the Late Noachian or
Early Hesperian such as those in Holden, Eberswalde, or Jezero craters may have been formed in conditions that had already become
less than favorable for life. Even though such sedimentary sites
provide invaluable information about surface hydrology and have
the advantage of clear stratigraphic context, their deposition in
the Late Noachian or Early Hesperian may have occurred on a sur-
face subject to a radiation environment that was similar to that of
Mars today. If the presence of the magnetic field was a necessary
requirement for habitability, and exploring habitable conditions
is the goal, this would imply that locations with more ancient
materials may give us the best hope for detecting traces of life from
early Mars.
Acknowledgments
We thank Jay Dickson, Bethany Ehlmann, and Ian GarrickBethell for helpful discussions. Reviews by Bob Craddock and an
anonymous reviewer improved the final manuscript. We gratefully
acknowledge financial assistance from NASA in support of coinvestigator participation on the ESA Mars Express High Resolution
Stereo Camera Team (JPL Contract 1237163).
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