Chronology of compressional deformation on Mars: evidence for a single and global origin

Planetary and Space Science 48 (2000) 1201–1211

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Chronology of compressional deformation on Mars: evidence for a single and global origin N. Mangolda; b;∗ , P. Allemandc , P.G. Thomasa , G. Vidala a Laboratoire

de Sciences de la Terre, ENS Lyon et UCBL, UMR CNRS 8515, 46, allÃee d’Italie, F-69364 Lyon, France b Orsay-Terre, UMR 8616, Bat.509, Univ. Paris-Sud, F-91405 Orsay, France c Laboratoire de Sciences de la Terre, UMR CNRS 8515, Univ. Claude Bernard, F-69622 Villeurbanne, France Received 25 May 1999; received in revised form 10 May 2000; accepted 17 May 2000

Abstract The chronology of compressional deformation on Mars is determined using chrono-stratigraphic relations and crater counting in order to provide new evidences for the origin of this deformation. Chronological relations between compressional tectonic features with dierent directions establish that a single deformation phase occurred in each studied region. On other hand, the intersections between tectonic features and craters permit to compare the relative age of the deformation in dierent regions. The results demonstrate that there may be a long lapse of time between the accumulation of Hesperian volcanic plains and the compressional deformation in each region. The possible age deduced for the compressional deformation on both Hesperian ridged plains and Noachian primitive terrains is Late Hesperian. A single and global phase may explain the compressional deformation because it is restricted to a single epoch. A global contraction of the planet may explain many properties of compressional tectonism on Mars according to the thermal evolution of the planets. Such process c 2000 does not exclude secondary sources of stress like Tharsis bulge that would control the geometry of structures at regional scales. Elsevier Science Ltd. All rights reserved.

1. Introduction Thrust faulting related to compressional deformation exists on each telluric planet. On Earth, such deformation is clearly controlled by plate tectonics. The origin of thrust faulting on one-plate planets like Mars, the Moon or Mercury is controversial. Regional processes like cooling of volcanic plains and global processes like global contraction have been invoked (Strom et al., 1975; Solomon and Chaiken, 1976; Spudis and Guest, 1988; Thomas et al., 1988). The case of Mars is especially complex because of its long geological evolution and the abundance of contractional features. On Mars, the most abundant compressional tectonic features are called wrinkle ridges in reference to contractional features of the Moon. Wrinkle ridges usually exist in the volcanic plains of the late Noachian and early Hesperian epochs which are referred

∗ Correspondence address: Orsay-Terre, UMR 8616, Bat.509, Univ. Paris-Sud, F-91405 Orsay, France. Tel.: 33-169154903; fax: 33-160191446. E-mail address: [email protected] (N. Mangold).

to as ridged plains (e.g. Banerdt et al., 1992). Features referred to as lobate scarps and high-relief ridges appear mostly in the primitive terrains of Noachian highlands. They are usually attributed to compressional deformation (Chicarro et al., 1985; Watters, 1993). The compressional origin of these structures is demonstrated by observations at contact between the Noachian highlands and volcanic plains. Lobate scarps sometimes cross the boundary and extend in ridged plains with a shape of a wrinkle ridge (Watters, 1993; Mangold et al., 1998). Several studies have examined the classi cation, distribution, spacing and widths of contractional features on the whole Mars (Chicarro et al., 1985; Banerdt et al., 1992; Watters, 1991). These tectonic features are observed on units of dierent ages. Regions devoid of contractional features are mainly Amazonian volcanic plains of the northern hemisphere. The center of Tharsis bulge is also devoid of ridges because this region is mostly Amazonian and furthermore it is submitted to a broad extensional stress eld (Banerdt et al., 1992). The anks of Tharsis bulge exhibit circumferential ridges indicating the in uence of the bulge on their formation. Various explanations attribute the origin of concentric ridges to the

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formation of the bulge (Banerdt et al., 1982; Tanaka et al., 1991; Zuber, 1995). Other ridged plains in the opposite hemisphere may not be related to Tharsis formation (Watters, 1993). Regional hypotheses like cooling and subsidence of volcanic plains are proposed to explain these ridged plains according to studies of the lunar mare ridges (e.g. Quaide, 1965). The thermal history of telluric planets like Mars suggests a slow secular cooling over geological times (e.g. Schubert et al., 1992). A period of global contraction is predicted by such cooling and a global ridge system is often cited as an evidence for this global contraction (Chicarro et al., 1985; Banerdt et al., 1992). The origin of Martian ridges is then ambiguous. Did local, regional or global processes form the ridges? Did these processes occur jointly? This study provides new arguments to discuss the problem of the origin of compressional tectonic features. These arguments are deduced from the chronology of deformation. Previous studies dated the deformation by counting all craters of the deformed terrain. This method did not give a precise age for deformation. For example, Watters (1993) found a peak of compressional deformation at early Hesperian, only “if tectonic features are roughly the same age as the units in which they occur”. Nevertheless, some stratigraphic relations near Valles Marineris attest of a time lapse between the volcanic accumulation and the formation of the ridges (Lucchita and Klockenbrink, 1981). The goal of this study is then to improve the chronology of compressional deformation. We especially use intersections between craters and tectonic features to estimate the time lapse between the lling and deformation of volcanic plains. Our results will show that compressional features are mostly Late Hesperian. This single age for all structures suggests a single and global origin. We propose that the global contraction due to thermal cooling is a possible solution for the origin of the deformation. 2. Crosscutting relations between compressive structures Most Martian regions only display a network of parallel ridges like the ridged plains surrounding Tharsis bulge. These regions were deformed under a single-direction compressive stress and probably a single deformation event. There are no evidences of polyphased deformations like for grabens on Tharsis bulge which are formed in at least six phases (Tanaka and Davis, 1988). However, several regions display cross-striking tectonic features which may be interpreted as resulting from dierent directions of stress. These dierent directions of stress may represent independent episodes of deformation as we can observe on Earth in mountain building. This section focuses on regions where intersections between tectonic features are common in order to establish the chronology of deformation.

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Fig. 1. (a) Viking Image 610a46 (23 S; 53 W). Part of Coprates Planum ridged plain. (b) Structural interpretation: Ridges R are orthogonal to the main compressive stress and ridges T are transpressive structures.

2.1. Coprates Planum Two or three directions are sometimes observed in the East of Coprates Planum (Fig. 1a). Some of them are not concentric to Tharsis bulge and are referred to as non-orthogonal ridges by Maxwell (1982). In Fig. 1b wrinkle ridges strike N–S (R1–R6), NE–SW (T1) and NW–SE (T2–T3). The orientation of ridges R corresponds to the general orientation of the concentric ridges around Tharsis. The direction of the main principal stress is then orthogonal to the set of faults R. Non-orthogonal ridges, like T1, are composed of several sections en-echelon. Such shapes are characteristics of transpressive ”push-up” structures. We observe that the ridges T relay ridges R. Indeed, T1 relays R3 from R6, T2 relays R2 from R3 and T3 is a relay between R2 and R4. The two sets of faults T1 and T2=T3 are conjugate faults with opposite sense of movement consistent with the direction of the main principal stress. Furthermore, the directions of ridges T are not random but correspond to the directions of the old fractures ◦ of the basement as observed on Viking image n 610a44 and discussed by Mangold et al. (1998). Indeed, wrinkle ridges are often observed in resurfaced terrains like volcanic plains but the primitive crust underlying these volcanic materials may have been aected by tectonic deformation early in the planetary evolution (e.g. Watters, 1993). The

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ambiguous geometry of their intersections. Ridges H and L are almost orthogonal, thus the direction of the main principal stress was not unique like on Coprates Planum. On the contrary, these observations and interpretations are in favor of a multidirectional compressive stress eld. This means that the two main directions of stress 1 and 2 were horizontal and of the same magnitude and that 3 was vertical. Although unusual in terrestrial plate tectonics, such stress eld would be the cause of all groups of ridges of Hesperia Planum at the same time. The occurrence of two main directions may correspond to a localization along old fracturations of the crust as proposed by Raitala (1988). 2.3. Consequences of observations ◦

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Fig. 2. Viking Image 418s33 (24 S; 243 W). Part of Hesperia Planum. Intersections between ridges H and L show that they formed at the same time.

fact that wrinkle ridges can form above underlying heterogeneities like faults and craters has been reported before (Maxwell, 1982; Raitala, 1988; Watters, 1993). The shape of ridges T1 is therefore a combination of compression and strike-slip while the direction of the main principal stress is orthogonal to ridges R. The interpretation of ridges T as a combination of strike-slip and compression was also deduced in others studies (Schultz, 1989). Ridges T are then due to the reactivation of faults that are non-orthogonal to the main direction of stress. Finally, these interpretations imply that the region of Fig. 1 was only aected by a single compressional phase under a single stress eld. We cannot identify more than one compressional stress eld in the hemisphere of Tharsis bulge regardless of whether ridges are concentric or not to the bulge. 2.2. Hesperia Planum Other regions of the planet show cross-striking ridges like Syrtis Major Planum, Hesperia Planum and Malea Planum. We only examine the case of Hesperia Planum because it displays a dense network of ridges but the conclusion that we propose is also valid for other regions. The two major trends of Hesperia ridges are NW–SE and NE–SW (Raitala, 1988) which correspond to ridges H and L in Fig. 2, respectively. Fundamental observations are done at the intersections between tectonic features. Ridge L1 is interrupted by ridges H2 and H3. The ridge H2 also interrupts ridges L2 and L3. Such geometry implies that ridges H are older than ridges L. On the other hand H2, is deformed by ridge L4. H1 does not crosscut L3. These last observations would imply that ridges H are younger than ridges L on contrary to the rst interpretation. Consequently, these two groups of structures have been formed necessarily at the same time to explain the

The comparison of the two previous examples demonstrates two important facts: (1) The stress eld creating ridges can be unidirectional or multidirectional and (2) in each region independently the ridges have been formed during the same deformation phase. In this conclusion, we do not take into account isolated ridges that may not be related to regional deformation but local processes like radial ridges on volcanoes. Indeed, such ridges are clearly due to subsidence of the volcano like on Alba Patera (Cattermole, 1990). The second fact pointing towards single stress elds and deformation phases is veri ed in every regions we focused on the planet: Lunae Planum, Coprates Planum, Hesperia Planum, Syrtis Major Planum, Arcadia Planitia, Cimmeria Terra and Arabia Terra. No observation has shown evidences for two or more distinct compressional phases. There is then no possibility to establish the chronology of deformation using intersections between tectonic features. However, the comparison of the relative age of deformation in each region is possible. If the cause of compressional deformation is dierent from one region to an other this deformation would most probably not be contemporary. On the contrary, if the origin was global, the age of the deformation would be the same in each region. Therefore, dating the deformation in each region may allow to discriminate between a regional and a global origin. 3. Chronology of deformation using crosscutting relations between craters and compressional features 3.1. The method of relative age determination The relative age of tectonic deformation is established following three principles: (1) crater density, (2) stratigraphic relations between deformed and undeformed terrains and (3) chronological relations between craters and structures. The two rst principles are classical methods for the determination of relative ages of planetary surfaces. The relative age of

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a given terrain corresponds to a density of craters of a given size. This method allows to date an area but not a group of structures, which are obviously younger than the terrain they aect. Stratigraphic relations are used when young terrains like volcanic plains bury one or several structures of the same region. This method gives an interval of time limited by (1) the age of the youngest terrain aected by deformation and (2) the age of the oldest terrain not aected by deformation. Unfortunately such an interval is often very approximate. Thus many studies assume that tectonic features are of the same age as the units in which they occur (Watters, 1993). We try to improve this method using the third principle, which is based on chronological relations between craters and structures: When a structure intersects a crater, this structure is younger than the crater. This method is described in Fig. 3. The studied area is homogeneous, so craters can be used to date this area. Before the deformation, 4 craters affect the studied area (Fig. 3A). When the deformation phase occurs the two craters 1 and 2 are aected by deformation (Fig. 3B). Until now the terrain is aected by other 3 craters and by a younger volcanic plain which buries partially one structure. We can deduce from that chronology that the age of the studied terrain is “7 craters”. We know from sketch A that 4 craters are older than the deformation. However, at present we can only conclude that craters 1 and 2 are older than the deformation because they are intersected. The maximum age of the deformation is then “5 craters” because it corresponds to the subtraction between the total number of craters (7 craters) and the number of intersected craters (2 craters). This is a maximum age because we cannot know if some of the undeformed craters are older than the deformation or not. On the other hand, we can de ne a minimum age for the deformation which corresponds to the younger terrain that buries the structure. Note that the crater counting must be translated in crater density and corresponds to a cumulative number of craters of a minimum size. The interval of the ages obtained is then restricted to the interval given (1) by the maximum age deduced from the total crater density and the density of transected craters and (2) by the minimum age deduced from stratigraphic relations with younger terrains. The determination of the lower age limit requires the knowledge of the chronological relations between craters and tectonic features. Fig. 4 shows an example of craters that are crosscut by ridges in Hesperia Planum. The intersection with the big crater (A) is obvious but the intersection with the smaller craters (B and C) needs detailed study of their rims. It is not always easy to decide if one crater is older (or not) than the crater it seems to intersect. Young craters can obliterate a structure, so all crater–structure relations have not been included in the count for that reason. High-resolution Viking images were used when possible in order to discriminate between the two alternatives. Allemand and Thomas (1995) described each possible relation between craters and structures. The large

Fig. 3. Schematic representation of chronological relations between craters, structures and layering. A: Before deformation, the studied area is aected by cratering. B: Deformation phase. Two craters are intersected by wrinkle ridges. C: At present, new craters formed and a younger terrain buries one tectonic feature. The age of the deformed terrain is N = 7. The maximum deformation age is N = 5 because of the 2 intersected craters. The minimum deformation age is limited by the age of the younger unit in the top right corner of sketch C.

number of crater–ridge intersections is not a coincidence because ridges are often initiated by heterogeneities in the subsurface layers (Allemand and Thomas, 1995; Mangold et al., 1998). This method is ecient to date compressional structures because these structures do not crosscut craters by chance. Thus, the number of craters transected by ridges is greater than the number of ridges aected by fresh younger craters because this last process is purely random. The determination of the upper age limit requires the study of chrono-stratigraphic relations between ridged terrains and younger terrains. The crater density of these younger terrains is established using geological maps of Mars (Scott and Tanaka, 1986; Greeley and Guest, 1987). The ages of these

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Fig. 4. Viking Image 365s69 (21 S; 242 W) in Hesperia Planum. The large crater A is transected by a large wrinkle ridges. Smaller tectonic features aect the rims of craters B and C. The 3 craters are older than deformation.

younger terrains are often imprecise due to the small surface they cover. Furthermore, several ridges may only be partially buried by younger volcanic llings, thus their ages would not correspond to the age of the last volcanic event. 3.2. Age of deformation in Hesperian ridged plains Fig. 5 is a map of the craters larger than 5 km and of the ridges in Hesperia Planum. The surface considered is S = 0:39 × 106 km2 . The number of intersections between ridges and craters is 22 with respects to a total of 76 craters larger than 5 km (Table 1). This crater density is normalized to a surface area of 1 × 106 km2 in order to enable comparison with the usual time scale. This density is usually called N (5) with a unit in craters=106 km2 (e.g.√Tanaka, 1986). The standard deviation on the density is ( N )=S where N is the number of craters. The relative age found for the volcanic plain is then N (5) = 192 ± 22 corresponding to the bottom of the Hesperian epoch. Indeed, this epoch is de ned by a density N (5) between 200 and 67 (Greeley and Guest, 1987). The density of craters that are not intersected gives a lower limit to the age of deformation at N (5) = 138 ± 19 in the middle Hesperian (Table 1). The upper limit is given by volcanic plains at the south of Hesperia Planum which are dated at about N (5) = 50 (Early Amazonian). Out ow channels south of Hadriarca Patera are also younger than ridges but their crater density is dif cult to constrain. The resulting interval of relative ages is then mainly restricted to the Late Hesperian. The plain of Syrtis Major Planum is a young ridged plain with N (5) = 124 ± 16 craters, thus at the beginning of Late Hesperian. This age is a little older than the age proposed on the map of Scott and Tanaka (1986), of N (5)=100±15, but younger than N (5)=165 proposed by Tanaka (1986). Using

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the ridged plain age we determined that the 11 transected craters give a lower limit of N (5) = 100 ± 15 craters for the compressional deformation. Volcanic terrains of Isidis Planitita give the upper limit. The border of Isidis is upper Amazonian (unit Aps) with at most 67 craters and the center of Isidis may be slightly older with N (5) = 75 ± 15 (Scott and Tanaka, 1986). The stratigraphy of Arcadia Planitia located at the West of Tharsis is more dicult to establish. The age of the studied area is xed at N (5) = 134 ± 19, thus at the end of the Early Hesperian. The few intersected craters give a lower age limit of the deformation at N (5) = 110 ± 17. The rst undeformed terrains are Amazonian plains (units Aph and Ael1) with a maximum age of N (5) = 67. Some of these terrains seem to be aected by ridges. However, this observation can be explained because the thickness of fresh lavas is supposed to be only 200 m as deduced from partially buried craters and knobs (Pike and Davis, 1984; De Hon, 1988). The result is that lavas do not completely bury tectonic features. The studied part of Lunae Planum displays 16 craters that give a maximum age of N (5) = 125 ± 15 for deformation. The studied area is surrounded by Valles Marineris and out ow channels. Valles Marineris walls are also younger than ridges but their age of about N (5) = 80 is not precise due to the small size of the area. Stratigraphic relations between ridges and out ow channels are not always clear. Some ridges appear on the oor of out ows in Chryse Planitia, but ow lines are often deviated by the topography of the ridges. Thus, most ridges may be older than out ow channels (Tanaka, 1986). Maja Valles at the East of the studied area seems undeformed by ridges, it then represents an upper limit for the deformation with an age of N (5) = 75 ± 20 (Tanaka, 1997). North of Lunae Planum, Chryse Planitia shows dierent ages between Lower Hesperian and Lower Amazonian (Tanaka, 1997), but the chronological relation with ridges is dicult to establish because some terrains display ridges and others not. Coprates Planum was studied by Allemand and Thomas (1995), who found N (5) = 141 craters for the lower limit. The upper limit is given by the nearby terrain of Sina Planum with an age of N (5) = 75 ± 10. Finally, the compressional deformations in ridged plains are always clearly younger than the formation of the volcanic plains they aect (Table 1). 3.3. Age of deformation in Noachian terrains Two Noachian regions have been studied. The case of Cimmeria Terra is detailed in Fig. 6a and b. The method using crater–structure intersections gives 19 intersections relevant to a total of 67 craters. The maximum age of deformation is then N (5) = 324 ± 47 in comparison to the mean age

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Fig. 5. Interpretative map of ridges and craters in Hesperia Planum (MC 22SE). Numbers indicate craters older than the compressive deformation. Table 1 Results of the chronology of compressional deformation in seven regionsa Name Image number

Latitude=longitude of area center

N ¿ 5 km

Inters.

Surface (MKm2 )

D ¿ 5 km

Max. age

Min. age

Hesperia Planum 25S=243 76 22 0.39 192 ± 22 138 ± 19 ∼ 50 MC22SE Coprates Planumb 20S=70 180 39 1.0 180 141 75 ± 10 MC18SW Lunae Planum 8N=68 78 13 0.52 150 ± 17 125 ± 15 75 ± 20 MC10SE–10SW Syrtis Major Planum 7N=288 56 11 0.45 124 ± 16 100 ± 15 75 ± 15 MC13SE Arcadia Planitia 25N=181 51 9 0.38 134 ± 19 110 ± 17 60 ± 10 MC8NW–15NE Cimmeria Terra 45N=200 67 19 0.15 450 ± 55 ∼ 166c — Vik 370s43– 47 Arabia Terra 28N=318 48 15 0.16 300 ± 40 ∼ 175c — MC12NE–13NW–5SC a N is the number of craters with diameter ¿ 5 km. Inters. corresponds to the number of these craters older than the structures. D is the crater density per 106 km2 . Max. age is the crater density established using crater–structure intersections and Min. age is the crater density of younger terrains. b Data from Allemand and Thomas (1995). c No meaningful standard deviation.

of the terrain of N (5) = 450 ± 55. This result does not provide a precise interval because of the lack of stratigraphic relationship with younger terrains. However, a better estimate results from the morphology of the craters. Indeed, the Noachian climate would have been hot and wet with a high degradation of the relief (e.g. Carr, 1996). In comparison, Hesperian climate was colder with low erosion rates. The Noachian craters can be distinguished from the Hesperian craters because their rim and central peak are degraded or non-existent (Craddock and Maxwell, 1993). There is also a gap of craters smaller than 20 km due to that degradation.

Most of the Hesperian craters can then be distinguished from the Noachian craters because they are not degraded and they are often surrounded by uidized ejecta. In Fig. 6 we see that 5 of the 19 craters intersected by structures may correspond to Hesperian craters. This result gives a maximum age of about N (5) = 166 taking the beginning of Hesperian to be N (5) = 200. The lack of degradation of the ridges themselves means that they formed after the Noachian–Hesperian transition in agreement with the previous argument. The second region studied in Noachian terrains is Arabia Terra (Table 1). The intervals given by the

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Fig. 6. (a) Mosaic of Viking Images 370s43–370s47 (45 S; 200 W) in Noachian terrains of Cimmeria Terra. (b) Interpretative map of craters and structures. N are Noachian craters aected by the deformation. H are craters older than the deformation that are supposed to be of Hesperian age.

total number of intersections is also not useful but 3 craters among the 15 intersected craters may correspond to Hesperian non-degraded craters. The maximum age may then be about N (5) = 175. No stratigraphic relation provides a lower limit for the deformations in these regions. Thus, the deformation in the Noachian terrains studied happened after the onset of the Hesperian but no more precise chronology can be deduced. 3.4. Interpretation Fig. 7 synthesizes results in both, the Hesperian and the Noachian terrains. Two main points are highlighted in this gure: (1) The chronology of compressional deformation in ridged plains (limited by the arrows) gives an age younger than the volcanic plains inside which they take place (with F in Fig. 7). This result implies that the deformation is not directly due to the lling and cooling of the plains. Our chronology moves forward the usual period of compressional deformation of the Early Hesperian (Watters, 1993) to the Late Hesperian. (2) The seven age intervals are all consistent with a deformation phase at the same time. Indeed, no region shows an age disconnected from the other regions. For example, no deformation stage is restricted to the Early Hesperian for a given region. Thus, a single deformation phase is possible between the crater

Fig. 7. Synthesis of age intervals (shown by arrows) for the seven regions studied. Crater densities are measured by counting craters with diameters ¿ 5 km according to the usual time scale of Mars (Tanaka et al., 1986). F represents the ages of the volcanic plains.

densities of N (5) = 100 and N (5) = 75 in the Late Hesperian (Fig. 7). This interval would correspond to an absolute age in the range of 3.7–3.6 Gyr in the scale of Neukum and Wise (1976) or 3.0 –2.1 Gyr in the scale of Hartmann et al. (1981). This deformation phase is thus restricted to a relatively short and late period. This

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chronology provides strong evidence for a single and large compressional phase on the whole planet. 4. Discussion 4.1. Possible origin of deformations Regional processes of compression are not consistent with a chronology requiring a single deformation phase on the whole planet. Furthermore, the regional processes most often quoted are cooling and subsidence of volcanic plains (e.g. Watters, 1993). Compressional stress would be a direct consequence of the formation of the plains by volcanism. Thus, this process can not explain the deformation if several hundreds million years separate the formation of the plains from their deformation as our chronology suggests (Fig. 7). One regional process that can produce global eect is the formation of Tharsis bulge. Ridges concentric to Tharsis are usually interpreted as a consequence of the formation of the bulge requiring isostatic adjustment or

exural loading (Sleep and Phillips, 1979; Banerdt et al., 1982). However, models do not reproduce all observations like the chronology of deformation. We refer in this matter to Watters (1993), who discussed the discrepancies between observations and models of stress around Tharsis bulge. Three points are focused on for the discussion: (1) Wrinkle ridges concentric to Tharsis are only Late Hesperian contrary to the grabens that are formed at least in six phases of deformation from Late Noachian to Amazonian (e.g. Tanaka and Davis, 1988). There is, therefore, no direct chronological relation between radial extension on the bulge and circumferential compression on its anks. (2) The distribution of contractional features shows that structures in the Eastern hemisphere are not oriented following the stress created by Tharsis bulge. Tharsis seems to have had no eect on these structures (Watters, 1993). (3) Models predict that compressional stress may decrease with distance from Tharsis but no gradient of deformation has been detected around Tharsis (Plescia, 1991). We then conclude, in agreement with Watters (1993), that Tharsis formation plays an obvious role in the orientation of ridges in the Western Hemisphere but that Tharsis does not explain a global compressional deformation on the whole planet. Another possible origin for contractional features could be global planetary processes like planetary despinning, lithospheric reorientation and global contraction (Banerdt et al., 1992). We can dismiss processes like planetary despinning and lithospheric reorientation because they would lead to the occurrence of conjugated networks of strike-slip and normal faults that are not observed (Melosh, 1977; Melosh, 1980; Banerdt et al., 1992). Global contraction was proposed in previous work in order to explain the global ridges system (Solomon and Chaiken, 1976;

Chicarro et al., 1985; Watters, 1993). This process is also involved to explain the planetary network of contractional features on Mercury and the Moon (Strom et al., 1975; Solomon and Chaiken, 1976; Spudis and Guest, 1988). Global secular contraction is due to the cooling of the interior of the planet which, however, is supposed to be a gradual process (e.g. Schubert et al., 1992). Global contraction could then explain compressional deformation while secondary processes like Tharsis may also play a role. Nevertheless, global contraction due to secular cooling should not be consistent with two observations: (1) Uniform cooling of an initial hot sphere with homogeneous and isotropic properties would not generate an oriented deviatoric stress at the surface (Watters, 1993). Thus, contractional features would be expected to occur in random or disorganized patterns rather than with the uniform orientations often observed. (2) Secular cooling is supposed to be a gradual process throughout the history of the planet but the observed chronology requires a peak of deformation and not a gradual deformation. A more precise description of global contraction is then needed to explain these discrepancies. 4.2. Orientations of tectonic features produced by global contraction Compressional tectonic features on Mars that may be due to planetary contraction display uniform orientations rather than the random orientations expected for thermal stress. Indeed, orientations are supposed to be random in the case of a uniform and isotropic sphere (Banerdt et al., 1992; Watters, 1993). However, Mars does not correspond to an isotropic sphere. For example, preferential directions of deformation can be due to the reactivation of buried structures or heterogeneities. Furthermore, if a non-hydrostatic stress eld derived from a regional sources is added to the hydrostatic thermal stress it would induce organized patterns of structures even if the magnitude of this stress is lower than the global stress (Watters, 1993). The formation of Tharsis bulge may have induced compressional stress around the bulge (e.g. Banerdt, 1982). If we assume that this stress was not high enough to break the crust and produce surface structures, this stress is only accommodated by elastic strain in the crust. Thus, if another stress is added like global compression, the strength of the crust may then be exceeded, and surface tectonic features can be formed. The consequence is that contractional features appear with orientations corresponding to secondary stresses around Tharsis. Such scenario can explain why ridges on Tharsis anks have an age dierent from the extensional deformation on the bulge. Such scenario would also explain the multidirectional stress observed in Hesperia Planum, which may have

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been devoid of residual stress when contraction occurred. The two main directions of ridges may have been the result of reactivation of buried structures (Raitala, 1988). Finally, contractional features are mostly observed in volcanic plains. We propose that the deformation is accommodated rst in these regions because the underlying crust is thinner and weaker than in the highlands. The organization of contractional features on the planets is then not an argument against global cooling. 4.3. Peak of compression produced by secular cooling In this section, we investigate scenarios for global contraction that could be consistent with a peak of compressional deformation. Global contraction is usually attributed to the thermal cooling of the interior. Thermal stress at the surface due to temperature change can be described by the relation (Turcotte, 1983; Solomon, 1986) t = (
Ti −
Tl )E=3(1 − );

(1)

where
Ti and
Tl are temperature changes in the interior (mantle+core) and in the elastic lithosphere, respectively,  is the Poisson coecient of the crust, E is the Young modulus of the crust and  is the volume coecient of thermal expansion. This stress is compressive only if the temperature changes in the lithosphere are smaller than the temperature changes in mantle and core. The change of the planetary radius due to that cooling can be calculated by (Banerdt et al., 1992) (
R=R) = t (1 − )=E;

(2)

where
R is the change of planetary radius R. If the temperature change in the lithosphere is neglected, a variation of 100 K of the interior would result in an isotropic compressional stress of about 150 MPa in the lithosphere. The global contraction would then be about 3.5 km. This contraction is not a process beginning with the formation of the planet. On one hand, stress is tensional when the change of temperature in the elastic lithosphere is higher than in the mantle. Turcotte (1983) predicts that the Moon may have sustained tensional stress during at least the rst billion years. Such tensile stress would have dominated the early Mars lithosphere during some hundred of million years (Schultz, 1985). On the other hand, the early evolution is dominated by the formation of the crust and by the changes of mineral phases in the mantle (Schubert et al., 1992). Calculations of expansion and contraction taking into account crustal dierentiation show that expansion has occurred during the rst two or three hundred million years (Schubert et al., 1992). After this time, contraction begins with rates of about 0.05 – 0.1% (1.5 –3 km) per billion years depending on crustal dierentiation (Fig. 15, Schubert et al., 1992). These results show that global contraction is not a primitive process, i.e. thermal compressive stress begins after the crust is formed. Elastic motions in the crust then rstly accommodate this stress. It may create tectonic features only if

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this stress reaches the maximal strength of the brittle crust. We can then compare the order of magnitude of the thermal stress and the crustal strength in order to understand if contraction can create compressive structures at the surface. Thermal stress can be deduced from relation (2) and the calculation of the radius change by thermal eects from the results of Schubert et al. (1992). Rates of radius change of about 1.5 –3 km=Gyr would then correspond to rates of thermal stress of about 150 –300 MPa=Gyr. If this stress is not accommodated it can reach 1 GPa after 4 Gyr. On the other hand, the maximum strength of the brittle crust corresponds to the strength at the brittle–ductile transition. This transition is controlled by the viscous ow of the lower crust that is related to composition, temperature, strain rate and pressure (e.g. Ranalli, 1995). The depth of this transition depends on the thermal gradient at a given time. We assume a thickness for the brittle crust of about 10 –15 km with a large thermal gradient that is supposed to have existed 3.5 Gyr ago. In absence of crustal cohesion, the strength is then given by the following relation (Ranalli, 1995):  = ÿgz with ÿ = 3 in compression. Density  is 2800 kg m−1 and gravity g = 3:72 m s−2 . The strengths  calculated are then in the order of 300 – 400 MPa for z = 10–15 km brittle crust. These values are of the same order of magnitude than the thermal stress after 1 or 2 Gyr. A global deformation phase can then occur in response of the global contraction late after the onset of cooling. 4.4. Peak of compressional tectonism produced by rapid contraction In the previous section we assumed that the global contraction is a gradual process but we observed that the consequence of such process is not obviously gradual. However, global contraction itself can result from a nongradual process. For example, Watters (1993) proposed that a peak of contraction occurred as a consequence of the volcanic pulse of the Early Hesperian. This hypothesis is attractive because it may explain a single and global episode of deformation creating structures in the volcanic plains as in the highlands. However, the pulse of volcanism is not restricted in time but spread over the Early Hesperian (see F in Fig. 7). Furthermore, tectonic features may be Late Hesperian so there is a large time gap between volcanic accumulation and deformation. This gap is shorter than 100 Myr in the Neukum and Wise (1976) chronology and longer than 500 Myr according to Hartmann et al. (1981). Such scenario is only possible in the rst chronology assuming that the cooling of the mantle anomalies inducing the volcanism takes some tens of million years, like on Earth (Turcotte and Schubert, 1982). On the other hand, the Martian core contributes to about one-fourth to thermal cooling (Schubert et al., 1992). A solid

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inner core may form depending on the sulfur concentration (Schubert and Spohn, 1990). The volume change due to the transition from liquid to solid iron is very large. The complete solidi cation of the core at a given time in Martian history is possible according to theoretical models (Schubert and Spohn, 1990). This solidi cation may then induce a “pulse” of global contraction and a high thermal stress in the crust. This hypothesis presently cannot be veri ed because of the poor knowledge of the evolution of Martian core. Nevertheless, this process shows that a global contraction peak due to internal processes is possible. Finally, several processes implying global contraction can produce a peak of compressive deformation on the surface of Mars. 5. Conclusion The chronology of the compression shows that (1) there is a long delay between ridged plains accumulation in the Early Hesperian and their deformation in the Late Hesperian and (2) all studied areas have deformation ages consistent with a single deformation phase in the Late Hesperian. These results show that compressive deformation is not related to regional processes like cooling of volcanic plains but may be the result of a single global deformation. The global contraction created by the thermal cooling of the Martian interior is able to explain such single, global and late deformation phase. Secondary stress elds, like that due to Tharsis bulge, are necessary to explain the orientation of structures. The arguments developed in the discussion are subject of the large uncertainties of the early Mars evolution, and global contraction is not a well-known process. The formation of large heterogeneities like the dichotomy and the Tharsis bulge probably created major changes in the thermal parameters of early Martian interior. All these processes may be connected: The expression of deformation at the surface of the planet only gives a partial view of this complex evolution. New topographic and imagery data from Mars Global Surveyor may provide new constraints on the role of these global processes on the development of the tectonic features observed on Mars. Acknowledgements We are grateful to Pr. Olivier Merle and an anonymous reviewer for their constructive comments. This research was supported by the French Programme National de Planetologie of the Institut National des Sciences de l’Univers (INSU-CNRS). References Allemand, P., Thomas, P.G., 1995. Localization of Martian ridges by impact craters: Mechanical and chronological implications. J. Geophys. Res. 100 (E2), 3251–3262.

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