The mineralogy of late stage lunar volcanism as observed by the Moon Mineralogy Mapper on Chandrayaan-1

Authors


Abstract

[1] The last major phases of lunar volcanism produced spectrally unique high-titanium basalts on the western nearside of the Moon. The Moon Mineralogy Mapper (M3) on Chandrayaan-1 has provided detailed measurements of these basalts at spatial and spectral resolutions necessary for mineralogical interpretation and mapping of distinct compositional units. The M3 imaging spectrometer acquired data in 85 spectral bands from ∼430 to 3000 nm at 140 to 280 m/pixel in its global mapping mode during the first half of 2009. Reflectance data of several key sites in the western maria were also acquired at higher spatial and spectral resolutions using M3's target mode, prior to the end of the Chandrayaan-1 mission. These new observations confirm that both fresh craters and mare soils within the western high-Ti basalts display strong 1 μm and weak 2 μm absorptions consistent with olivine-rich basaltic compositions. The inferred abundance of olivine is observed to correlate with stratigraphic sequence across different mare regions and absolute ages. The apparent stratigraphic evolution and Fe-rich compositions of these basalts as a whole suggest an origin from evolved residual melts rather than through the assimilation of more primitive olivine-rich sources. Mare deposits with spectral properties similar to these late stage high-Ti basalts appear to be very limited outside the Procellarum-Imbrium region of the Moon and, where present, appear to occur as small areas of late stage regional volcanism. Detailed analyses of these new data and supporting measurements are in progress to provide further constraints on the mineralogy, olivine abundance, and compositions of these final products of lunar volcanism and the nature and evolution of their source regions.

1. Introduction

[2] The history of volcanism on the western nearside of the Moon is unique in both its duration and the basaltic compositions it left exposed on the lunar surface. Together, Oceanus Procellarum and Mare Imbrium compose the largest regional expanse of the lunar maria, with deposits ranging from the early history of mare volcanism through its last major phases [e.g., Hiesinger et al., 2003]. Various remote sensing techniques have dated many of the flows in these regions as younger than 3.0 Ga [Boyce, 1976; Schaber, 1973a; Bugiolacchi and Guest, 2008] with some areas having been dated as young as ∼1.2 Ga [Hiesinger et al., 2003] and <1 Ga [Schultz et al., 1976; Schultz and Spudis, 1983]. Such flows are believed to represent the final products of mare volcanism emplaced almost 2 Gyr after the youngest basalts collected by sample return missions.

[3] The last major phases of lunar volcanism cover large areas of Oceanus Procellarum and Mare Imbrium as well as smaller regions within several surrounding maria [Hiesinger et al., 2000, 2003]. These flows have high inferred titanium contents due to their relatively blue ultraviolet to visible reflectance properties [e.g., Charette et al., 1974; Pieters, 1978; Johnson et al., 1991; Blewett et al., 1997] and orbital elemental measurements [Davis, 1980; Elphic et al., 2000, 2002]. Telescopic studies first identified a relatively strong 1 μm feature and a weaker 2 μm absorption associated with these mare regions, suggesting the presence of a significant component of olivine or Fe-rich glass in western high-Ti basalts [Pieters et al., 1980]. Subsequent studies of Clementine data supported interpretations of abundant olivine and high-FeO contents within these basalts, but lacked the spectral resolution for more detailed mineralogy [Staid and Pieters, 2001; Lucey, 2004]. In the current study, visible to near-infrared reflectance data acquired by the Moon Mineralogy Mapper (M3) on Chandrayaan-1 are used to investigate the mineralogy of these late stage basalts and implications for the heterogeneity and evolution of their source regions.

2. Previous Work

[4] Among the lunar maria, the compositional properties and evolution of the late stage high-Ti basalts in Oceanus Procellarum and Mare Imbrium appear to be unique. These flows include the last major phases of lunar volcanism [Schultz et al., 1976, Schultz and Spudis, 1983; Young, 1977; Hiesinger et al., 2000, 2003] and occur in a distinct compositional region of the Moon known as the “Procellarum KREEP Terrain” [Jolliff et al., 2000; Wieczorek and Phillips, 2000; Haskin et al., 2000]. Extensive studies of the western maria have characterized their unique compositions, geologic setting and ages. Some key studies are discussed here as background; however, more extensive discussions of previous work are summarized in other publications [Head, 1976; Whitford-Stark and Head, 1980; Basaltic Volcanism Study Project, 1981; Wilhelms, 1987; Head and Wilson, 1992; Jolliff et al., 2006].

2.1. The Distribution and Diversity of Mare Basalts

[5] The distribution and volumes of lunar volcanic deposits provide a record of the thermal evolution of the Moon providing a model for the fundamental processes of crustal evolution elsewhere in the solar system [Taylor, 1989; Head and Wilson, 1992]. The Apollo and Luna missions sampled a subset of the Moon's basalts, erupted between ∼4.3 and 3.1 Gyr ago [Lofgren et al., 1981; Taylor et al., 1983; Dasch et al., 1987]. The wide range of TiO2 contents within the returned samples has been used as a chemical property for separating them into groups that include high-, low-, and very low Ti basalts which generally cannot be related to the same source regions [Papike et al., 1998; Longhi, 1992]. High-Ti basalts in the sample collection include Apollo 11 low- and high-K samples, as well as Apollo 17 basalts from the eastern nearside of the Moon ranging in age from 3.55 to 3.85 Ga (a recent summary of age dating studies is given by Stöffler et al. [2006]).

[6] Earth based and orbital geochemical measurements have greatly expanded what is known about the history and diversity of lunar volcanism beyond the regions sampled by the Apollo and Luna missions. Telescopic studies first classified nearside basalts into a number of spectral types based on continuum slope, albedo, and near-infrared absorption features of mature mare soils [e.g., Pieters and McCord, 1976; Johnson et al., 1977; Pieters, 1978]. Apollo X-ray and gamma-ray data also provided compositional information for limited areas of the Moon [e.g., Davis, 1980]. In subsequent years, additional knowledge about the global distribution of the lunar basalt compositions resulted from the Galileo SSI flybys of the Moon [Belton et al., 1992; Greeley et al., 1993; Pieters et al., 1993], Clementine UVVIS and NIR data [Nozette et al., 1994; Pieters et al., 1994] and Lunar Prospector elemental data [Binder, 1998; Lawrence et al., 1998; Elphic et al., 1998]. Recently, a return to the Moon composed of a wide range of missions equipped with advanced remote sensing instruments (Smart-1, Kaguya/SELENE, Chang'e-1, Chandrayaan-1 and LRO) is providing the information necessary for a much more detailed view of the geologic history of lunar volcanism [e.g., Foing et al., 2006; Haruyama et al., 2009; Ono et al., 2009; Goswami and Annadurai, 2009; Pieters et al., 2009; Besse et al., 2011; Whitten et al., 2011].

[7] The observed distribution of lunar basalt types suggests that mare volcanism was regionally complex with no simple correlation between composition and absolute age [e.g., Pieters, 1978; Giguere et al., 2000; Hiesinger et al., 2003, 2011]. Though a wide range of basalt types exist within many nearside basins, some broad compositional trends are apparent. Early volcanism consisted of a range of titanium and iron contents (and therefore densities) in the nearside basins after the period of heavy bombardment. Within these early deposits, high-Ti deposits are most prevalent earlier in the stratigraphy of the eastern nearside [e.g., Hiesinger et al., 2000]. Outside of the major basins, basalts emplaced in areas of thicker crust [Zuber et al., 1994] were less diverse and are dominated by high-albedo spectrally red low-Ti deposits [e.g., Pieters, 1978; Wilhelms, 1987]. For these basalts, basin controlled crustal thickness and magma density likely played a significant role in emplacement volumes and/or composition of extruded lavas [Solomon, 1975; Wilhelms, 1987; Head and Wilson, 1992; Wieczorek and Phillips, 1998] by controlling magma ascent, timing of crystal fractionation and settling, and possible assimilation of feldspathic materials or other crustal materials [cf. Neal and Taylor, 1992; Lofgren et al., 1981]. The last major phases of lunar volcanism, occurring on the western nearside, once again produced high-Ti basalts distinct in composition from earlier lunar volcanism.

2.2. The Western High-Titanium Basalts

[8] The dark volcanic flows exposed on the western nearside are believed to have TiO2 abundances similar to the older Apollo 11 and 17 high-Ti basalts based on their ultraviolet to visible spectral properties [e.g., Pieters, 1978; Whitford-Stark and Head, 1980] and orbital geochemical data [e.g., Davis, 1980; Elphic et al., 2002; Prettyman et al., 2006]. The general distribution of these basalts within Oceanus Procellarum and Mare Imbrium can be seen in the outlined regions of Figures 1a and 1b (the western high-Ti basalts show up as blue-green in Figure 1b). These last major eruptions also occurred in areas associated with unusually high abundances of radiogenic elements [Soderblom et al., 1977; Adams et al., 1981; Lawrence et al., 1998] within the Th-rich province, observed by the Lunar Prospector gamma-ray spectrometer known as the “Procellarum KREEP Terrain” [Jolliff et al., 2000]. A component of this Th enrichment is believed to be associated directly with these maria, further supporting the presence of unique mare sources with adequate heat for the eruption of these late stage lunar basalts [e.g., Soderblom et al., 1977; Jolliff et al., 2001; Flor et al., 2003].

Figure 1.

Late stage lunar volcanism on the western nearside appears as a distinct red hue in this M3 color composite. (a and b) Western high-titanium basalts appear dark and spectrally blue in Galileo SSI 0.56 μm albedo and standard color ratio images (0.41/0.56 μm, 0.76/0.99 μm, 0.56/0.41 μm). (c) M3 integrated band strength color composite in which these western basalts are red due to a strong 1 μm integrated band depth and weak 2 μm band (red for 1 μm IBD, green for 2 μm IBD, blue for reflectance at 1.58 μm). (d) A schematic of the M3 IBD parameters used in Figure 1c. Dashed white lines in Figures 1a–1c indicate location of the study area shown in Figure 2.

[9] Though unsampled by landed missions, the ages of the western high-Ti basalts are constrained by a wide range of stratigraphic and crater-dating studies. These dark and relatively blue volcanic flows cover older mare units across the eastern portion of Oceanus Procellarum and much of the Imbrium basin. Within Mare Imbrium, the basaltic flows flood small and large Eratosthenian craters and cover older Imbrian deposits, providing a context for their stratigraphic age [Wilhelms, 1987]. Flow fronts within the Mare Imbrium deposits can be seen in low-Sun Apollo orbital photographs and individual stratigraphic sequences of these basalts were mapped by Schaber [1973b]. In northern Oceanus Procellarum, some areas of the high-Ti flows have been dated as young as twice the age of Copernicus, or 1.6–2.0 Ga [Young, 1977] and even Copernican in age (<1 Ga) [Schultz and Spudis, 1983]. More recent crater counts of large areas of Oceanus Procellarum and Mare Imbrium by Hiesinger et al. [2000, 2003] and for Mare Imbrium by Bugiolacchi and Guest [2008] together date exposed high-Ti volcanism on the western nearside from >3 Ga to as recently as 1.1 Ga.

[10] The late stage western basalts are spectrally distinct from the older, high-Ti compositions on the eastern nearside of the Moon. Though the ultraviolet to visible reflectance properties of these basalts are similar to those sampled by Apollo 11, Pieters [1978] classified the western maria into separate spectral classes based on differences in the strength and shape of the mafic band near 1 μm. Since telescopic spectra of both mare soils share similar weak 2 μm bands, but younger western maria were observed to have a stronger and broader absorption centered near 1 μm, Pieters et al. [1980] interpreted these basalts to have an additional ferrous absorption from olivine and/or iron-bearing glass.

[11] Subsequent studies of these basalts using higher spatial resolution Clementine UVVIS and NIR data [Staid and Pieters, 2001; Staid et al., 2002] examined fresh mare craters and associated soils within the high-Ti Procellarum deposits and Imbrium flows. These studies determined that the strong and asymmetric 1 μm absorptions within the mare soils and relatively weak 2 μm ferrous absorptions were also present for crystalline materials excavated from varying depths throughout these flows by optically immature craters. Based on the strength and pervasive nature of the asymmetric and strong 1 μm absorption, Staid and Pieters [2001] concluded that the presence of abundant olivine within the emplaced basalts was the most likely explanation for the unique spectral properties of these basalts. This study also concluded that the young high-Ti basalts were very iron rich (>20 wt % FeO) based on a comparison of mare craters and soils from the western high-Ti deposits to the basalts sampled by Apollo 11 in Tranquillitatis. Examination of individual flow sequences of these basalts in Mare Imbrium [Schaber, 1973b] further suggested that olivine content was increasing in subsequent eruptions within those flows [Staid and Pieters, 2001]. Other studies of Clementine data based on the FeO-mapping approach of Lucey also predicted high-FeO contents for the western high-Ti basalts (∼18–22% for broad regions); however, these Clementine FeO estimates did not show large differences in iron content from earlier Tranquillitatis high-Ti basalts [Lucey et al., 1998, 2000].

[12] Iron mapping from Lunar Prospector independently identified the western mare regions as very iron rich, with large areas having iron abundances of 22–23% FeO and some areas as high as 25%. These FeO estimates for the western mare soils were significantly larger than those observed for the older eastern high-Ti maria (∼16–20%), and the authors reasoned that the western basalts likely represent the most iron-rich large expanses of mare basalts on the Moon [Elphic et al., 2002]. Differences between Prospector and Clementine FeO mapping results for the relative iron content of the older and younger high-Ti basalts may result from regional variations in mineralogy, such as olivine or ilmenite content that are not fully captured in the Clementine iron-mapping algorithm [Lucey et al., 2000; Staid and Pieters, 2000; Clark and McFadden, 2000; Elphic et al., 2002; Chevrel et al., 2002]. Since Prospector measured the FeO content of average mare soils (all samples of which contain significant quantities of highland contamination [Laul and Papike, 1980; Haskin and Warren, 1991]) rather than actual basalts, Elphic et al. [2002] concluded that the mare basalts of Procellarum may have higher-FeO contents than any known lunar samples. Alternatively, these authors noted that the younger age of the western basalts could result in less average-highland contamination of the mare soils.

[13] More recent compositional analyses of the western high-Ti basalts include a quantitative analysis of combined UVVIS and NIR Clementine data by Lucey and Steutel [2003] and Lucey [2004]. These studies used a radiative transfer model to examine the global distribution of lunar mineralogy within the Clementine data and predicted average abundances of ∼50% olivine by volume within the western high-Ti basalts [Lucey, 2004]. Additional studies of mare craters in the western basalts, over expanded wavelength range of the UVVIS and NIR camera, as well as ROLO telescopic data, have also supported interpretations of abundant olivine and iron-rich compositions within these late stage basalts [e.g., Staid et al., 2004; Staid and Stone, 2007]. Clementine UVVIS analyses and crater counts were also integrated, in a regional study of Mare Imbrium, by Bugiolacchi and Guest [2008], who observed that the igneous petrogenesis of basalts in this region appear to have evolved through time to more TiO2 and FeO-rich melts. The unique compositional properties of the western high-Ti basalts indicate significant changes in the nature of lunar volcanism over time. Though lunar volcanism was regionally complex during the peak of basalt emplacement, the Moon's last large-scale eruptions appear to have produced unique compositions not observed in earlier basalts.

2.3. Spectral Analysis of Mare Craters and Soils

[14] Because space weathering and mixing alter the reflectance properties of the lunar surface over time, it is difficult to characterize the mineralogy of emplaced basalts from remote measurements of optically mature soils. In contrast, relatively crystalline and unweathered basaltic regoliths associated with young impact craters retain diagnostic absorption features that can be interpreted based on measurements of returned lunar samples [e.g., Pieters, 1977; McCord et al., 1981]. However, previous spectral data obtained by telescopes have lacked the spatial resolution to observe small craters excavating individual basaltic units [Pieters, 1993]. The western high-Ti basalts are particularly thin, and many craters resolvable with telescopic data are believed to have excavated earlier low-Ti flows [Pieters, 1977; Pieters et al., 1980].

[15] Recent data from the M3 instrument on Chandrayaan-1 provides the detailed spectral resolution of Earth-based telescopic data at spatial resolutions of 70–140 m/pixel. Studies of Clementine UVVIS and NIR data (at ∼100–500 m/pixel) demonstrate that these resolutions are adequate to resolve the spectral properties of immature crater deposits sampling individual volcanic flows [e.g., Staid and Pieters, 1996, 2000; Kramer et al., 2008a, 2008b]. Unlike previous orbital or Earth based observations, M3 provided both the high spatial and spectral resolutions capable of investigating the detailed reflectance properties of small mare craters. These observations allow for a more direct characterization of basalt mineralogy than measurements of lunar soils that have been altered by space weathering and nonmare contamination.

3. The Moon Mineralogy Mapper: Instrument and Analysis

3.1. The M3 Instrument and Lunar Mapping Data

[16] The M3 imaging spectrometer was a guest instrument on India's Chandrayaan-1 mission which launched on 22 October 2008 and mapped the lunar surface from a polar orbit through August 2009. M3 acquired visible to infrared reflectance data at spatial and spectral resolutions capable of measuring discrete basaltic flows within the lunar maria. Most of the M3 data were collected in a global mapping mode that covered the wavelength range of ∼430 to 3000 nm in 85 spectral bands at 140 to 280 m/pixel spatial resolutions. Small amounts of data were also acquired over targeted regions at the full spectral and spatial capability of M3 (259 spectral bands and spatial resolutions of ∼70 m). The mission ended partway through its nominal 2 year mapping period in late August 2009, after a loss of communications with the satellite. Despite an abbreviated mission, M3 was able to cover more than 95% of the Moon in its global mode of operations [Boardman et al., 2011]. By comparison, only a small number of target mode images were collected at the full spatial and spectral resolution of the instrument. Several of these early targeted data collections were located in the western high-Ti basalts. Though M3 was designed to operate over four 3 month optical periods at solar illumination angles of 30° or less, thermal issues experienced by the spacecraft required operation well outside of the planned viewing geometries of the instrument. As a result of these thermal conditions and the limited duration of mapping, much of the M3 global coverage and existing targeted data were acquired under relatively high incidence angles (i > 45°) outside of the nominal viewing geometries planned for instrument mapping [Boardman et al., 2011]. In May 2009, the Chandrayaan-1 orbit was raised from 100 km to 200 km, resulting in global mode data after this time period being collected at a reduced 240 m/pixel spatial resolution. Additional details of the operation aspects of the M3 instrument during lunar mapping are given by Boardman et al. [2011], and details of the M3 instrument design and capabilities are presented by R. O. Green et al. (The Moon Mineralogy Mapper (M3) imaging spectrometer for lunar science: Instrument, calibration, and on-orbit measurement performance, submitted to Journal of Geophysical Research, 2011).

3.2. M3 Data Calibrations

[17] The performance and radiometric calibrations of the M3 imaging spectrometer are described in detail by Green et al. (submitted manuscript, 2011). In its full resolution mode (target mode), M3 measured the spectral range from 406 nm to 2991 nm with ∼10 nm sampling at 12 bit digitization. In global mode, 2 × 2 spatial averaging and 2 to 4 times spectral averaging was applied over predetermined wavelength ranges to reduce data downlink volumes. M3 data used for this study were processed through a series of steps that include dark signal subtraction, mapping and replacement of bad detector elements, application of laboratory-determined flat fields, secondary on-orbit flat-field adjustments, and calibration to radiance (Green et al., submitted manuscript, 2011). Refinements of the mission radiometric calibrations are ongoing, and various levels of calibration have been defined as they are iterated. All data presented here are based on the ‘K’ radiance calibrations of M3. Known issues in this version of the calibration that have been addressed in subsequent calibrations include a scattered light component below 1 μm and the presence of an electronic panel ‘ghost’ (Green et al., submitted manuscript, 2011). For this study, the M3 K data were converted to apparent reflectance by normalizing the radiance data by the calculated solar flux at the time of acquisition based on

equation image

where L is the M3 measured radiance, F is the calculated solar flux based on a solar curve corrected for Sun-Moon distance, and i is the solar incidence angle during acquisition. After calibration to apparent reflectance, a small “ground truth” correction was applied to all M3 apparent reflectance data. This correction is a fixed 85 band spectrum derived as a global correction for the 85 band global mode data to smooth residual band to band noise based on the smoothly varying spectral properties of lunar sample measurements. The correction used here is based on a fit to the global M3 data set using the Empirical Flat Field Optimal Reflectance Transformation (EFFORT) spectral polishing approach of Boardman [1998]. Since a similar correction has not yet been derived for the small subset of 259-band target mode data, a higher-order polynomial local EFFORT correction was applied to the targeted data presented in section 4.2. Unless otherwise stated, reflectance data presented in this paper have not been corrected to a standard viewing geometry for comparisons presented in this work. Instead, efforts have been made to compare spectra across a limited range of similar phase angles for mineralogic interpretation. The thermal corrections of Clark et al. [2011] have been applied to the global mosaics and targeted data presented in this paper, but not to individual global mode orbits examined in this study. As a result, presentations of global mode spectra are truncated at 2.5 μm, beyond which additive contributions due to emissivity become most significant.

3.3. M3 Data Analysis

[18] To perform preliminary compositional assessments, global mosaics were created by reducing the spatial resolution of the M3 data by a factor of 10 [Boardman et al., 2011]. After calibration of the data to apparent reflectance and the application of thermal corrections [Clark et al., 2011; Green et al., submitted manuscript, 2011], a range of spectral parameters was calculated to explore broad mineralogical differences across the Moon. Figure 1c shows a color composite of a nearside M3 mosaic composed of spectral parameters that distinguish mare regions based on their 1 and 2 μm integrated band depths (IBD) and near-infrared albedo. A schematic that depicts how these parameters are calculated is shown in Figure 1d: the total integrated 1 μm band area between 789 and 1308 nm is displayed in red, the total integrated 2 μm band area between 1658 and 2498 nm is shown in green, whereas the reflectance at 1580 nm is shown in blue. The M3 science team determined that this combination of parameters best captures the first-order variance in lunar spectral properties. In this combination, highlands with weak mafic absorptions are blue, whereas regions rich in mafic minerals are yellow/green to orange/red, depending on relative ferrous band strengths resulting from differences in mineralogy and optical maturity. The late stage western maria are clearly distinguished from earlier basalts in this color composite due to their relatively weak 2 μm absorptions and strong 1 μm ferrous bands (red hues in this color composite image).

[19] To characterize the reflectance properties of the spectral groups observed in the M3 IBD mosaics, global data strips were identified within the M3 data set that cross-mare unit boundaries at moderate solar-illumination angles. The majority of M3 data of the western basalts was acquired under relatively high illumination and phase angle conditions (mostly >55° phase for large regions of Imbrium and northern Procellarum) due to operational constraints resulting from thermal problems experienced by the Chandrayaan-1 spacecraft during lunar mapping [Boardman et al., 2011]. These high-phase observations result in highly shadowed and illuminated surfaces along topographic slopes, and therefore, complicate the spectral characterization of immature materials associated with mare craters. During early mapping, however, several orbits of global mode data were acquired under more moderate phase angle conditions (<45° phase), which included mare unit boundaries identified as areas 1, 2, and 3 in Figure 2. These orbits (and corresponding regions) were selected for initial characterization of the spectral properties of the western basalt units mapped in the low spatial resolution IBD mosaics.

Figure 2.

Subset of the spectral parameter image shown in Figure 1 centered over the mare deposits of northern Oceanus Procellarum and Mare Imbrium. The extent of Figure 3 is outlined as the white rectangle, and the approximate locations where the reflectance properties of fresh mare craters were sampled from different spectral units (Figure 4) are indicated with numbers. Letters refer to named craters on the Moon: L, Lichtenberg; A, Aristarchus; C, Copernicus; E, Euler; M, Marius.

[20] Some of the youngest age estimates for the western basalts are associated with the high-Ti units in northwestern Procellarum, near the 20 km wide Lichtenberg crater. Schultz and Spudis [1983] interpreted the youngest mare basalts within this region to embay upon the ejecta of this potentially Copernican aged crater that retains visible ejecta rays. Based on these stratigraphic associations, an age estimate of <1 Ga was suggested by Schultz and Spudis. Crater age dating has produced somewhat older, but also relatively young, age estimates of ∼1.7 Ga [Hiesinger et al., 2003] and 2.2 Ga [Morota et al., 2008]. Full spectral and spatial resolution data of Lichtenberg were acquired during a brief period of targeted collections that occurred over the western nearside of the Moon prior to loss of Chandrayaan-1. The M3 targeted data of Lichtenberg have been processed to apparent reflectance in the same manner as the global mode data and were further thermally corrected based on the approach of Clark et al. [2011]. M3 data of the crater and surrounding regions were examined in spectral parameter images, and spectra of mare units surrounding Lichtenberg were compared after sampling fresh mare craters identified in M3 images. Additional discussion of the processing associated with the targeted data and the results of mare unit comparisons are presented in section 4.3.

4. Results

4.1. M3 Mosaics and Spectral Parameter Images

[21] A subset of the global IBD mosaic centered on northern Oceanus Procellarum and Mare Imbrium is shown in Figure 2. Several large and contiguous areas within the maria are observed to be relatively homogeneous within this M3 spectral parameter. Mare basalts that predate the western high-Ti flows have a yellow hue in this color composite due to strong ferrous absorptions at both 1 and 2 μm. Highland regions are primarily blue due to higher near-infrared reflectances and lower mafic band strengths; however, localized variations in mafic band strength are observed in association with fresh craters (greenish blue in the composite image). White outlines in Figure 2 provide the approximate boundaries of the late stage high-Ti basalts within this region. These deposits are distinct from earlier basalts and range from light red to dark red within the IBD mosaic. Within the high-Ti basalts, large regions appear as coherent groups based on their integrated 1 and 2 μm band strengths. These groupings are roughly mapped in Figure 2, by further separating the high-Ti basalts into two broad spectral groups based on the relative strengths of their 1 and 2 μm mafic absorptions (dashed versus dotted lines). More subtle spectral variations occur southwest and southeast of Aristarchus that are not distinguished in Figure 2. These areas may represent additional spectral and stratigraphic boundaries within the high-Ti basalts that are not captured by the two broader spectral groups.

[22] Comparison with crater age estimates for the mare basalts [e.g., Hiesinger et al., 2003] shows strong associations between the basalts with the strongest 1 μm versus 2 μm integrated band depths (dark red in Figures 1c and 2) and the youngest-dated basalts. However, these age units appear to be less contiguous than the spectral groupings observed in the M3 IBD parameter image. The mare age maps of Hiesinger et al. [2000, 2003] were based on representative areas within spectral units defined by differences in Clementine and Galileo color ratio images (based on UV/VIS and 1 μm band depth ratios). In addition to distinguishing mare compositional units, the Clementine and Galileo ratio composites are sensitive to spectral variations resulting from local mixing with nonmare materials and variations in optical maturity. The M3 data presented in Figure 2 suggests that compositional boundaries for age dating could also be considered over the broader spectral and stratigraphic groups observed in the M3 IBD parameter mosaics. For example, in Mare Imbrium, the two spectral and stratigraphic groups seen in Figure 2 can be directly correlated with flow boundaries mapped by Schaber [1973b]. These flows were interpreted to originate from a source near Euler crater (E in Figure 2) and flow generally northeast across the Imbrium basin in distinct stratigraphic sequences [Schaber, 1973b]. Areas associated with the upper flows (phases II and III of Schaber [1973b]) have noticeably stronger 1 μm versus 2 μm integrated band depths that appear as coherent spectral units within these phases. By comparison, stratigraphically older high-Ti basalts (light red in Figure 2 and corresponding to phase I of Schaber [1973b]) have intermediate 1 μm versus 2 μm band strength differences that lie between the older pre-Eratosthenian low-Ti basalts and the final phases of volcanism in these regions. Similar trends can be seen in other regions of Imbrium and Procellarum in Figure 1c and 2, where the darkest red flows in these parameter images (weakest 2 μm versus 1 μm ferrous absorptions) appear to be the stratigraphically youngest units, overlying older high-Ti maria with less dramatic differences in 1 and 2 μm band strength.

[23] The reflectance properties of spectral groups observed in the M3 IBD mosaics are examined in more detail along several unit boundaries identified as areas 1, 2, and 3 in Figure 2. Global mode data of each type area were calibrated to apparent reflectance as described previously. Figure 3 shows an IBD mosaic of the full-resolution global mode data covering areas 1 and 2, as well as the general location of this region north of Aristarchus within Oceanus Procellarum. The mare boundary between areas 1 and 2 in Figure 3 is marked by a change in spectral properties from the older low-Ti deposits in the north to the younger high-Ti flows to the south. Mare soils to the north have a yellow hue in this composite due to strong ferrous absorptions at both 1 and 2 μm. Soils to the south are distinctly redder due to a strong 1 μm, but weak 2 μm integrated band strengths. Though this spectral boundary in the mare soils can be clearly seen in both images shown in Figure 3, it is more distinct in the lower spatial resolution mosaics that have been thermally corrected than it is in the higher-resolution global mode data that has not. This difference in the resulting IBD composite images is likely to result from the additive thermal emission signal's effect on the 2 μm IBD parameter within these low-albedo basalts.

Figure 3.

Global mode M3 data showing the locations of study areas 1 and 2 near a mare boundary in northern Oceanus Procellarum. (left) Location of the M3 study areas within a reduced resolution IBD mosaic of the nearside. (right) M3 IBD composite of areas 1 and 2 as observed in an individual orbit of M3 global mode data. Mare craters within the older low-titanium unit to the north appear as yellow in this composite while craters in the younger high-titanium unit to the south appear red.

[24] Due to the high spatial resolution of M3, it is also possible to directly observe the reflectance properties of small optically immature (or “fresh”) craters sampling these mare units. Such craters can be identified in this global mode IBD composite due to their increased 1 and 2 μm ferrous band strengths. Fresh mare craters within the young high-Ti deposits are distinctly redder (strong 1 μm band and weak 2 μm band) than the fresh craters within the older low-Ti unit to the north, which has strong ferrous absorptions at both 1 and 2 μm (bright yellow).

4.2. M3 Global Mode Spectral Data

[25] Small mare craters and associated ejecta with strong mafic bands were sampled from each study area and averaged to compare the spectral properties of relatively crystalline materials across compositional units. Resulting spectra from these mare units (areas 1, 2, and 3 in Figure 2) are compared before and after continuum removal in Figure 4. Continuums were removed from the global mode data by fitting a straight line to each spectrum from 730 to 1618 nm and dividing each spectrum by the resulting continuum. These spectra, based on the full-resolution global mode data, have not yet been thermally corrected and include an additional and increasing signal at longer wavelengths caused by thermal emission. As a result, apparent reflectance spectra presented in Figure 4 are limited to 2.5 μm, and analyses of continuum-removed spectra focus on wavelengths below 2 μm.

Figure 4.

(a and b) Spectra of fresh mare craters associated with the three distinct spectral groups identified in Figures 2 and 3. Low-Ti basalts that predate higher-Ti flows (e.g., area 1) are dominated by 1 and 2 μm pyroxene absorptions typical of most pre-Eratosthenian lunar basalts. The reflectance properties of craters within the younger high-titanium basalts (areas 2 and 3) have strong but longer-wavelength 1 μm absorptions and relatively weak 2 μm bands, consistent with abundant olivine within these basalts. (c) Laboratory reflectance spectra of lunar pyroxenes and olivine mineral separates (Brown University, NASA RELAB facility) [Pieters, 1993; Isaacson and Pieters, 2010].

[26] Spectra obtained from fresh craters in each of the three mare study areas are plotted together in Figures 4a and 4b. Moderate to large craters that appear to be excavating underlying units of different compositions were avoided based on differences in the spectral properties of their ejecta relative to smaller craters and surrounding soils. Laboratory reflectance measurements of lunar sample separates common to mare basalts are presented in Figure 4c for comparison to the M3 results [Pieters, 1993; Isaacson and Pieters, 2010].

[27] Low-Ti basalts that predate higher-Ti flows (area 1, Figure 4) display strong 1 and 2 μm pyroxene absorptions typical of most pre-Eratosthenian lunar basalts. The apparent position of the ferrous band center near 0.98 μm and 2 μm is consistent with typical high-Ca pyroxene compositions (Figure 4c) [Adams, 1974; Cloutis and Gaffey, 1991]. The specific band center of the 2 μm ferrous absorption in these and other global mode spectra presented in Figure 4 is less reliable than positions at 1 μm, due to a lack of thermal corrections in these data. Previous studies of Clementine data have shown that fresh mare craters retain spectral slope differences related to titanium content [Staid and Pieters, 2000] that have been demonstrated in laboratory studies of mature mare soils [Charette et al., 1974]. The UV/VIS properties of the older low-Ti units observed in area 1 are redder than craters sampled in areas 2 and 3, consistent with lower-Ti contents. The mare crater spectrum from area 1 is also brighter than areas 2 and 3, consistent with a lack of abundant opaques such as the Ti-rich phase ilmenite. However, since only a small number of craters were averaged from each unit, absolute albedos and band strengths cannot be reliably compared across study areas due to potential differences in the optical maturity of the craters sampled.

[28] M3 spectra collected from craters within the younger high-Ti basalts (areas 2 and 3) have strong, but longer wavelength 1 μm absorptions and relatively weak 2 μm bands. Olivine reflectance spectra display a broad and asymmetric composite absorption near 1 μm and lack a 2 μm absorption that is present in pyroxenes (Figure 4c) [Adams, 1975; Singer, 1981; Burns, 1993], and the M3 global mode spectra are consistent with previous observations that have attributed these spectral properties to the likely presence of olivine within these basalts [Pieters et al., 1980; Staid and Pieters, 2001; Lucey, 2004]. The UV/VIS ratio of these deposits are also relatively blue and dark compared to the older low-Ti deposits of area 1, consistent with their derivation from materials associated with the upper high-Ti rather than underlying low-Ti units. Area 3 also lies within the high-Ti western basalts but appears to be stratigraphically older than basalts to the north associated with area 2. Basalts in this region also have significantly older crater age estimates (∼2.8 Ga compared to 1.2 to 2.1 Ga for the spectral unit to the north sampled by area 2 [Hiesinger et al., 2003]). Mare soils in this region appear as light red in the M3 IBD color composites (e.g., Figures 1c and 2), with strong 1 μm ferrous bands and weaker 2 μm bands that are intermediate in strength between the youngest high-Ti basalts and pre-Eratosthenian low-Ti basalts. On average, craters sampled in this region lack 1 to 2 μm band strength differences as strong as for area 2 but still display weaker 2 μm bands than typical pyroxene-rich basalts. However, the spectral properties of craters in area 3 were observed to vary significantly from those with more typical pyroxene absorptions to craters with properties similar to those in area 2 (strong 1 μm but weak 2 μm absorption). Differences between the band positions and strength of the area 2 and area 3 basalts are further compared in the continuum-removed spectra presented in Figure 4b. In Figure 4b, area 2 basalts have a band center that lies beyond 1 μm and is intermediate in position between the pyroxene-rich basalts in area 1 and the olivine-rich basalts in area 2. Similarly, the 2 μm band is slightly stronger than it is for the area 2 basalts while the 1 μm bands have similar strengths in this continuum-removed plot.

4.3. M3 Targeted Data

[29] Targeted data of several areas of the western high-Ti basalts were acquired prior to loss of communication with Chandrayaan-1 satellite. These data have approximately 3 times the spectral and twice the spatial resolution of the global mode M3 data. Because these basalts are known to be relatively thin in many locations, targeted data may provide the best opportunity to examine small fresh craters from the youngest flows. Targeted data over Lichtenberg crater were acquired during the second optical period and were processed to apparent reflectance as described in section 3.2. These data were acquired at a higher illumination angle than the global mode data studied for areas 1–3, resulting in phase angles of ∼60° near Lichtenberg. In addition to a local high-order EFFORT correction used to smooth band-to-band noise [Boardman, 1998], the 0.705 μm channel was removed and interpolated with surrounding channels at 0.695 and 0.715 μm due to residual noise at that wavelength. Although calibrations of the targeted data are still preliminary, the spectral properties of this data complement the global-mode data results and provide additional spectral detail for mineral interpretations.

[30] M3 data of Lichtenberg crater and surrounding regions is presented in Figure 5 as an IBD color composite overlaid on a thermal channel at 2.78 μm (as brightness) to show mafic band variations in relation to the morphology of the crater and surrounding topography. Older low-Ti basalts can be seen to the southwest and north of Lichtenberg crater and appear yellow in this color composite image, particularly in the ejecta of small craters that have strong mafic absorptions at both 1 and 2 μm. The younger high-Ti unit appears red due to weaker 2 μm band strengths and can be seen on the eastern side of the image embaying older basalts and topographically higher regions east of Lichtenberg. In Figure 2, these flows can be seen as a contiguous spectral unit with the youngest stratigraphic units sampled in area 2 and units to the north dated as young as 1.3 Ga by Hiesinger et al. [2003]. Bright ejecta, presumably rich in feldspathic materials, can be seen as blue rays crossing the low-Ti units to the north of Lichtenberg and portions of the low-Ti units southwest of the crater. Small fresh craters within areas L1 and L2 were sampled and averaged together to obtain the reflectance spectra presented in Figure 6. Even relatively small craters (<1 km diameter) appear to be excavating through the high-Ti basaltic flows in some areas, as indicated by anomalously orange ejecta blankets compared to other small fresh craters in Figure 5. One of the most visible examples of a crater excavating an underlying unit is marked by a white arrow in Figure 5 and has a diameter of ∼600 m. Since depth of excavation scales with diameter (1:10 for simple craters) [Melosh, 1989], the thickness of the high-Ti flows are estimated to be less than ∼60 m in this area. In order to isolate the most diagnostic signature of a potential olivine component in these basalts and to avoid sampling the underlying unit, only locations within small craters that displayed the strongest 1 μm versus 2 μm ferrous band strengths were averaged for these comparisons (red craters in Figure 5). As a result, the reflectance properties presented are more likely to represent an end-member of potentially olivine-rich compositions within these basalts rather than the average spectral properties of all fresh small craters sampling the unit.

Figure 5.

IBD composite of M3 target mode data of Lichtenberg crater (20 km) and surrounding mare deposits. The IBD color composite has been overlaid on a thermal channel at 2.78 μm (as brightness) to show mafic band variations in relation to the morphology of the crater and surrounding topography. Older low-titanium basalts (L1) can be seen to the southwest and north of Lichtenberg crater and appear yellow in this color composite image. The younger high-titanium unit (L2) appears red due to weaker 2 μm band strengths and can be seen on the eastern side of the image embaying older basalts and topographically higher regions east of Lichtenberg. A white arrow locates a small crater that appears to be excavating through the thin high-titanium basalts to expose the underlying low-titanium deposits.

Figure 6.

M3 target mode reflectance spectra obtained from fresh craters in regions L1 and L2 south of Lichtenberg crater. Spectra in Figure 6b are presented after removal of a straight-line continuum from the reflectance properties of craters within each mare unit.

[31] Spectra obtained from several fresh mare craters within each unit are presented for comparison in Figure 6a. Figure 6b presents these each spectrum after normalizing the reflectance data to a straight-line continuum fit to visible and near-infrared peaks on either side of the 1 μm band. For the low-Ti units, the continuum was fit at 0.74 μm and 1.54 μm, while the continuum for the high-Ti basalts was fit to the 0.74 μm and 1.82 μm target mode channels.

[32] The crater spectra of the low-Ti unit southwest of Lichtenberg crater have strong absorptions near 1 and 2 μm that are consistent with high-Ca pyroxenes. Band centers for these absorptions appear to occur at ∼0.98 and 2.1 μm. Brightness and absolute band strength differences between the two groups of crater spectra cannot be reliably compared because only a small number of locations were sampled, resulting in potential differences in absolute maturity. In particular, only the smallest fresh craters could be sampled from within the high-Ti unit because of the thinness of those basalts, increasing the likelihood of including distal and more mature soils within the averaged spectra of those craters. Fresh craters from the low-Ti group do, however, have a redder UV/VIS slope consistent with their lower-Ti contents and would be expected to be brighter due to a relatively low amount of opaques compared to the high-Ti basalts.

[33] The high-Ti basalts in Figure 6 show spectral properties consistent with the global mode measurements of area 2, which forms a contiguous spectral unit with these basalts in the M3 IBD mosaics (Figures 2 and 3). The asymmetric shape of the strong 1 μm absorption has a distinct secondary feature between 1.2 and 1.3 μm, and the 2 μm absorption within these basalts is also relatively weak consistent with the presence of olivine. The 2 μm band occurs at even longer wavelengths than the lower-Ti, high-Ca pyroxene and may, in part, be attributable to chromite inclusions associated with the olivine [Cloutis et al., 2004; Isaacson and Pieters, 2010] or could be the result of the difficulty of thermally correcting data of craters that include topographic slopes.

5. Discussion

5.1. Mare Crater Spectra and Implications for Basalt Mineralogy

[34] The reflectance properties of both mare soils and fresh craters observed by M3 confirm the presence of a strong 1 μm absorption and weak 2 μm band within the western high-Ti basalts [Pieters et al., 1980; Staid and Pieters, 2001; Lucey, 2004]. The M3 data, however, provide a vastly improved combination of spectral and spatial resolution to directly observe the reflectance properties of small craters excavating individual compositional units within the maria. These observations allow for improved discrimination of potential components which share similarly positioned absorptions, such as olivine and iron-rich glass, and examination of specific stratigraphic sequences of basalts and their mineralogic associations.

[35] Olivine reflectance spectra display a broad and asymmetric composite absorption near 1 μm and lack a 2 μm absorption that is present in pyroxenes (Figure 4c) [Adams, 1975; Singer, 1980; Burns, 1993]. A central absorption just beyond 1 μm is caused by iron in the M2 site, and ‘wing’ absorptions near 0.9 and 1.3 μm result from iron in the M1 site [Burns, 1993]. In mineral mixtures with olivine, an asymmetric band near 1 μm, with a ‘secondary’ band near 1.3 μm, may be most visible [Singer, 1981; Pieters et al., 1980]. Within the global mode data examined, area 2 (youngest high-Ti unit north of Aristarchus in Figures 2 and 3) displays the broadest 1 μm absorption and weakest 2 versus 1 μm band strengths among the mare basalts. The 1 μm absorption associated with these basalts is much broader than in the neighboring pyroxene-rich basalts of area 1 and is centered beyond 1 μm. A longer-wavelength absorption centered between 1.2 and 1.3 μm is also apparent in this spectrum. The high-Ti basalts that include area 2 are, therefore, interpreted to contain the most visible and presumably highest component of olivine of the three mare regions studied in the global mode data. Another region with a strong 1 μm band but a weak 2 μm absorption in Figure 3 occurs in southeastern Aristarchus, a region previously interpreted as containing olivine [Le Mouélic et al., 1999]. This area is examined in more detail using M3 data in the companion paper by Mustard et al. [2011].

[36] Other common lunar minerals that produce absorptions beyond 1 μm and lack significant 2 μm features include plagioclase feldspar and Fe-rich glass. Plagioclase feldspar has weak absorptions at 1.2 μm and lunar glasses that have a broad absorption centered around 1 μm [Adams, 1975]. However, weak plagioclase absorptions are not expected to be visible in the presence of even small amounts of darker mafic minerals present in basalts [e.g., Crown and Pieters, 1987] and lose absorption features near 1 μm when shocked, such that the mineral has proven difficult to detect even in most highland regions. Fe-rich lunar glasses also contain broad absorptions extending beyond 1 μm; however, these absorptions are more symmetrical than the composite absorption observed within olivine. Fe-rich glasses produced by rapid cooling of basalts could contribute to broad absorptions observed in the weak spectral features associated with the western high-Ti soils [Pieters et al., 1980]. However, FeO-rich glasses are less likely to produce a strong and distinctly olivine-like signature in fresh craters excavating and mixing materials from depth, as observed in the M3 data of area 2.

[37] Olivine is less absorbing than pyroxenes and is thus likely to be masked by pyroxene absorptions, unless the olivine is present in relatively large abundance [Singer, 1981; Pieters et al., 1980; Mustard and Pieters, 1987]. The presence of the distinct olivine shape within spectra from area 2 suggests that either the olivine is very abundant relative to pyroxene (olivine/pyroxene > 1) or factors such as grain size and mineral associations within these basalts allow light to reflect more easily from the olivine-rich component. Since darkening components such as ilmenite may be associated with the pyroxene component or matrix of a basalt, independently of crystals of olivine, these basalts may have a lower olivine/pyroxene ratio than would be inferred by a linear interpretation of the strength of the olivine features observed in the M3 data.

[38] The high-Ti basalts sampled in area 3 lie west of Aristarchus and appear as a light red hue in the M3 IBD mosaic in Figure 2. The distribution of these basalts in both Procellarum and Mare Imbrium, as well as comparisons to mare age estimates, indicates that they generally predate the spectral unit sampled at area 2. These mare soils also have relatively weak 2 μm absorptions compared with older, surrounding low-Ti basalts, but the differences in band strength are less extreme than observed for the basalts sampled in area 2. The spectral properties of fresh mare craters sampled from area 3 (Figure 4), also display a broad, long-wavelength 1 μm band and weaker 2 μm absorptions consistent with the presence of some olivine. The spectral properties of the ferrous bands within the area 3 basalts appear intermediate between those of the pyroxene-rich, low-Ti basalts and the youngest high-Ti basalts. Area 3 high-Ti basalts are, therefore, interpreted to have at least some olivine present, but a lower average olivine/pyroxene ratio than the basalts sampled in area 2. Alternatively, differences in grain size and associations between olivine and the opaques (e.g., chromite, ilmenite) within these basalts could also result in the observed spectral differences from area 2 with similar olivine contents. Exposure to greater amounts of vertical mixing with the underlying and pyroxene-rich low-Ti basalts could also result in lower inferred olivine contents than the stratigraphically younger high-Ti basalts. However, the global data sampled at area 3 have a similar spectrally blue slope like area 2 and do not display an increase in albedo or reddening that would be expected from such mixing of mare materials.

[39] The targeted M3 data near Lichtenberg crater provides an opportunity to examine the shape of the 1 μm feature within the youngest high-Ti basalts in greater detail. Unlike the global mode spectra presented in Figure 4, these data were thermally corrected using the approach of Clark et al. [2011], which should provide improvements in the shape of the data near 2 μm, relative to the uncorrected data. These basalts are spectrally contiguous with area 2 in the global IBD parameter mosaic, but as described previously, are interpreted to be very thin in the region near Lichtenberg. The comparison in Figure 6 attempts to isolate some of the most olivine-rich basalts associated with craters in the high-Ti unit for comparison to the older, low-Ti and pyroxene-rich basalts to the west. It is again noted that the calibrations of the targeted data used for this study are very preliminary and a local smoothing of the data based on the EFFORT method was necessary to make these preliminary comparisons using the targeted data. The phase angles of the data themselves were quite high (∼60°), also making comparisons of samplings of crater materials complicated by shadowed and illuminated slopes. The resulting spectra for both the high- and low-Ti units, however, are consistent with craters sampled in the lower spatial and spectral resolution global data. In particular, the low-Ti basalts in this region have spectral properties that are consistent with laboratory and telescopic measurements of typical pyroxene-rich lunar basalts. The crater materials sampled from the high-Ti unit have spectral properties that are consistent with a high abundance of olivine. Due to the limited locations of fresh crater ejecta that could be sampled in this small region of targeted data, it is not possible to determine how typical these olivine rich materials are within other craters sampling this spectral unit.

[40] Previous studies of Lunar Prospector and Clementine data have characterized the late stage western, high-Ti basalts as among the most FeO-rich basalts on the Moon [Lawrence et al., 2002; Staid and Pieters, 2001]. Determining whether the olivine compositions of these basalts are also relatively FeO-rich is relevant to the evolution and source regions of these basalts. The shape and position of the 1 μm composite absorption for olivine are known to vary systematically from MgO- to FeO-rich compositions due to the position of Fe2+ in the crystal structure [Burns, 1993; Sunshine and Pieters, 1998]. The systematic changes in olivine band positions and shape with composition are well documented, and detailed modeling of olivine has been demonstrated in the laboratory and in remote sensing data [Burns, 1993; Sunshine and Pieters, 1998; Sunshine et al., 2007]. Qualitative comparisons of the spectra in Figure 6 to laboratory spectra of MgO-rich and FeO-rich (fosterite and fayalite) olivines indicate that the late stage, high-Ti basalts may be relatively FeO-rich. In particular, the longer-wavelength M2 absorption appears comparatively strong in these basalts relative to their central M2 absorption, producing an overall 1 μm band shape and long-wavelength edge more similar to the FeO-rich than MgO-rich end-members measured in laboratory studies [Sunshine and Pieters, 1998]. However, the modeling of absorption band positions in lunar basalts containing both olivine and pyroxene is a complex problem and beyond the current scope of this paper. Furthermore, grain size can also affect the shape of olivine spectra and complicate the identification of MgO-rich versus FeO-rich olivines [e.g., Lucey, 1998]. As a result, no conclusions about the composition of the olivine in these basalts can be reached without additional calibration and modeling of the M3 targeted data. The presence of high-FeO basalts can be expected to have lower Fo contents than normal but probably not lower than Fo50 in quantities detectable with remote sensing data.

5.2. Stratigraphy and Distribution of Olivine-Rich Basalts

[41] The inferred abundance of olivine in the western high-titanium basalts is observed to vary stratigraphically, with the uppermost flows in several areas (dashed lines, Figure 2) displaying the broadest 1 μm absorptions and weakest relative 2 μm band strengths. Stratigraphically older high-Ti flows within Imbrium and Procellarum (e.g., area 3 in Figure 2) also appear to contain olivine but at lower or more variable concentrations relative to their pyroxene abundances. This stratigraphic pattern is observed in a number of different regions with a wide range of estimated ages [Hiesinger et al., 2003, 2011]. Some of the youngest dated basalts near Lichtenberg crater [Schultz and Spudis, 1983; Hiesinger et al., 2003, 2011] are included in and contiguous with the most olivine-rich group, sampled in area 2 in global mode data and near Lichtenberg in the M3 targeted data. Additionally, the uppermost flows of basalts in some regions with older age estimates, such as in central Mare Imbrium and areas of southern Procellarum, also have similar spectral properties to these olivine-rich basalts. For example, the distribution of flows with increasing inferred olivine content are also a good match to phases of basalt emplacement mapped by Schaber et al. [1973] and support previous interpretations of increasing olivine abundance with subsequent eruptive phases of iron and titanium-rich basalts in this region [Staid and Pieters, 2001].

[42] The volcanic history of related basaltic units in the Marius Hills complex (bottom of Figure 2) is examined in more detail in a companion study of M3 data by Besse et al. [2011]. This study also identifies olivine-rich basalts within the youngest regional high-Ti flows in the crater Marius and surrounding deposits. Together, these new M3 observations suggest that the late stage basalts exhibit a pattern of increasing olivine abundance with subsequent emplacement that is widespread within the western high-Ti deposits, producing recognizable sequences of mare volcanism over more than a billion years of lunar history. Future studies of sequential phases of these basalts, such as those mapped by Schaber [1973a] in Mare Imbrium represent important sites for trying to substantiate and expand on these observations. For example, the apparent stratigraphic evolution and Fe-rich compositions of these basalts suggest an origin through evolved residual melts rather than through the assimilation of more primitive (Mg-rich) olivine-rich sources. If future studies confirm that these basalts have moderate to high-FeO contents that increase with subsequent eruptive phases, such basalts might originate from a mantle magma chamber that is undergoing mineral fractionation and crystal settling, resulting in increased FeO contents to the residual magma, which might be tapped at different times. This would result in increased FeO contents for subsequent basaltic olivines, and with concurrent decreases in abundances. Alternatively, if olivine FeO abundances are found to decrease with subsequent eruption, this may indicate that source regions are deepening over time or that a single olivine-rich source is being melted repeatedly.

[43] The late stage western basalts in Imbrium and Procellarum (which extend into northwestern Frigoris) represent the largest exposure of this unique spectral type that we have seen on the Moon. Preliminary examination of global-scale M3 parameter images does not reveal similar spectral properties among the other large-scale mare deposits on the Moon. However, small deposits with similar spectral characteristics can be identified in M3 parameter images within several areas outside of Mare Imbrium and Procellarum. Where present, these deposits appear to occur as small flows near or superimposed on older and more extensive maria or as isolated mare ponds. One such region occurs within small mare ponds southwest of Humorum where deposits that appear as nondescript mare ponds in Clementine color ratio images display strong 1 μm and weak 2 μm bands in M3 IBD images. Other small deposits that stand out in the M3 IBD parameter occur along the edges of Mare Frigoris. However, because the global M3 parameter images are nonunique, each potential basalt occurrence will require further examination using the full spatial and spectral resolution M3 data before they can be confirmed to represent basalts with similar spectral properties as those observed in this study.

5.3. Summary

[44] M3 observations of western nearside maria have provided new information about the mineralogy and emplacement history of these last major phases of lunar volcanism. The M3 data have confirmed previous observations that these basalts exhibit a unique combination of strong 1 μm and weak 2 μm ferrous bands [Pieters et al., 1980; Staid and Pieters, 2001; Lucey, 2004] and have observed these properties throughout large regions of soils and craters within these deposits. The improved spatial and spectral resolution of the M3 data allows direct observation of small craters sampling individual basaltic flows and improved discrimination of diagnostic absorption bands for the interpretation of mineralogy. Based on these new observations, we interpret the western basalts as having significant and variable quantities of olivine that can be directly observed in the reflectance properties of fresh mare crater regoliths. The abundance of olivine within these basalts appears to vary stratigraphically, with the uppermost flows being most olivine rich. This mineralogical trend appears to exist across regions and absolute ages of emplacement suggesting a common pattern of magma evolution of within these high-Ti basalts. Some small mare deposits with similar spectral properties may exist as the final products of mare volcanism in other isolated regions of the Moon; however, large regions of basalts with similar compositions (high titanium, high iron, and abundant olivine) are not observed as other major mare deposits elsewhere on the Moon.

[45] This initial study of the M3 data for basalts on the western nearside only begins to explore the wealth of information provided by these new imaging spectrometer measurements. The discrete stratigraphic sequences of the high-Ti basalts occurring in Mare Imbrium and portions of Procellarum will be important targets for future study. More detailed examination of M3 data within such previously mapped flows should provide information about mineralogical trends related to the evolution of mare basalt composition between eruptive phases as well as potential changes due to fractionation of materials prior to and during emplacement. Several concurrent studies of M3 data have demonstrated the utility of quantitative techniques such as the Modified Gaussian Method (MGM) to interpret olivine and pyroxene compositions from M3 data [e.g., Isaacson et al., 2011; Klima et al., 2011]. Future work will attempt to provide estimates of the relative abundance and composition of pyroxene and olivine to provide further constraints on their source regions, temporal evolution and emplacement mechanisms.

Acknowledgments

[46] M3 is funded as a mission of opportunity through the NASA Discovery program. The M3 science team and science validation are supported through NASA contract NNM05AB26C. We thank ISRO for the opportunity to fly as a guest instrument on Chandrayaan-1 and gratefully acknowledge their many contributions to the acquisition and return of the M3 data.

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