Brown University Planetary Geosciences - Lunar …Lunar ﬂoor-fractured craters as magmatic intrusions: Geometry, modes of emplacement, associated tectonic and volcanic features,

$Page 1: Brown University Planetary Geosciences - Lunar …Lunar ﬂoor-fractured craters as magmatic intrusions: Geometry, modes of emplacement, associated tectonic and volcanic features,$
Icarus 248 (2015) 424–447

Contents lists available at ScienceDirect

Icarus

journal homepage: www.elsevier .com/ locate/ icarus

Lunar floor-fractured craters as magmatic intrusions: Geometry, modesof emplacement, associated tectonic and volcanic features, andimplications for gravity anomalies

http://dx.doi.org/10.1016/j.icarus.2014.10.0520019-1035/� 2014 Elsevier Inc. All rights reserved.

⇑ Corresponding author.E-mail address: [email protected] (L.M. Jozwiak).

Lauren M. Jozwiak a,⇑, James W. Head a, Lionel Wilson b

a Department of Geological Sciences, Brown University, Providence, RI 02912, USAb Lancaster Environment Centre, Lancaster University, Lancaster LA1 4YQ, UK

a r t i c l e i n f o a b s t r a c t

Article history:Received 29 May 2014Revised 8 October 2014Accepted 12 October 2014Available online 15 November 2014

Keywords:VolcanismMoon, interiorMoon, surface

Lunar floor-fractured craters are a class of 170 lunar craters with anomalously shallow, fractured floors.Two end-member processes have been proposed for the floor formation: viscous relaxation, and subcrat-er magmatic intrusion and sill formation. Recent morphometric analysis with new Lunar ReconnaissanceOrbiter Laser Altimeter (LOLA) and image (LROC) data supports an origin related to shallow magmaticintrusion and uplift. We find that the distribution and characteristics of the FFC population correlatesstrongly with crustal thickness and the predicted frequency distribution of overpressurization valuesof magmatic dikes. For a typical nearside lunar crustal thickness, dikes with high overpressurization val-ues favor surface effusive eruptions, medium values favor intrusion and sill formation, and low valuesfavor formation of solidified dikes concentrated lower in the crust. We develop a model for this process,make predictions for the morphologic, morphometric, volcanic, and geophysical consequences of the pro-cess and then compare these predictions with the population of observed floor-fractured craters. In ourmodel, the process of magmatic intrusion and sill formation begins when a dike propagates verticallytowards the surface; as the dike encounters the underdense brecciated region beneath the crater, themagmatic driving pressure is insufficient to continue vertical propagation, but pressure in the stalled dikeexceeds the local lithostatic pressure. The dike then begins to propagate laterally forming a sill whichdoes not propagate past the crater floor region because increased overburden pressure from the craterwall and rim crest pinch off the dike at this boundary; the sill then continues to inflate, further raisingand fracturing the brittle crater floor. When the intrusion diameter to intrusion depth ratio is smallerthan a critical value, the intrusion assumes a laccolith shape with a domed central region. When the ratioexceeds a critical value, the intrusion concentrates bending primarily at the periphery, resulting in a flat,tabular intrusion. We predict that this process will result in concentric fractures over the region of great-est bending. This location is close to the crater wall in large, flat-floored craters, as observed in the craterHumboldt, and interior to the crater over the domed floor in smaller craters, as observed in the craterVitello. A variety of volcanic features are predicted to be associated with the solidification and degassingof the intrusion; these include: (1) surface lava flows associated with concentric fractures (e.g., in the cra-ter Humboldt); (2) vents with no associated pyroclastic material, from the deflation of under-pressurizedmagmatic foam (e.g., the crater Damoiseau); and (3) vents with associated pyroclastic deposits from vul-canian eruptions of highly pressurized magmatic foam (e.g., the crater Alphonsus). The intrusion of basal-tic magma beneath the crater is predicted to contribute a positive component to the Bouguer gravityanomaly; we assess the predicted Bouguer anomalies associated with FFCs and outline a process for theirfuture interpretation. We conclude that our proposed mechanism serves as a viable formation process forFFCs and accurately predicts numerous morphologic, morphometric, and geophysical features associatedwith FFCs. These predictions can be further tested using GRAIL (Gravity Recovery and Interior Laboratory)data.

� 2014 Elsevier Inc. All rights reserved.

http://crossmark.crossref.org/dialog/?doi=10.1016/j.icarus.2014.10.052&domain=pdf

http://dx.doi.org/10.1016/j.icarus.2014.10.052

mailto:[email protected]

http://dx.doi.org/10.1016/j.icarus.2014.10.052

http://www.sciencedirect.com/science/journal/00191035

http://www.elsevier.com/locate/icarus

L.M. Jozwiak et al. / Icarus 248 (2015) 424–447 425

1. Introduction

Floor-fractured craters (FFCs) are a class of 170 lunar craterscharacterized by their anomalously shallow, fracture-coveredfloors. First described in detail by Schultz (1976) with initial map-ping of >80 FFCs by Wilhelms (1987), these craters are divided intoeight morphologic subclasses (Schultz, 1976; Jozwiak et al., 2012).Each subclass possesses distinct morphologic characteristics,including features such as mare deposits, moat features, dark halodeposits, either convex-up or flat-floor profiles, and predominatelyradial, concentric, or polygonal fracture features.

There are two proposed formation mechanisms for floor-frac-tured craters (FFCs): (1) viscous relaxation, wherein the crater floorrebounds to fill the crater at a rate controlled by the subsurface vis-cosity structure resulting in an overall amplitude shallowing oflong-wavelength crater topography (Masursky, 1964; Daneš,1965; Cathles, 1975; Hall et al., 1981; Dombard and Gillis, 2001),and (2) magmatic intrusion and sill formation, wherein a dikepropagates from the mantle, stalls beneath the crater, then spreadslaterally beneath the crater and inflates, forming a laccolith andlifting up and fracturing the overlying crater floor (Brennan,1975; Schultz, 1976; Wichman and Schultz, 1996; Jozwiak et al.,2012).

Jozwiak et al. (2012) analyzed the morphology of 170 FFCsusing new Lunar Orbiter Laser Altimeter (LOLA) data and LunarReconnaissance Orbiter Camera (LROC) images to assess supportfor each of the proposed formation mechanisms. They found nodirect support for FFC formation via viscous relaxation; they did,

b

-3.5

-3

-2.5

-2

-1.5

-1

-0.5

0

0 5 10 15 20 25 30 35 40 45 50

Distance [km]

Ele

vati

on [

km]

Bohnenbergerc

Diameter= 33 km

-4

-3

-2

-1

0

1

2

0 20 40 60 80 100 120 140Distance [km]

Ele

vati

on [

km]

TychoDiameter= 86 km

a

Depth = 4.8 km

Fresh Depth = 3.0 kmBohnenberger Depth = 1.1 km

Fig. 1. Comparison of crater floor profiles including: (a) the fresh, unmodified crater Tyccrater Humboldt (27.2�S, 80.9�E), with a gently upbowed and shallowed floor profile, (c)crater floor region separated from the crater wall by a distinct ‘‘v’’-shaped trough, and (d)for (a, b, and c) are shown using LOLA 512 px/deg gridded data. The given crater depthscalculated according to Pike (1980). (d) This half-space model of a viscously relaxed crcharacteristic viscous relaxation time s/g) required to achieve a crater floor profile simi

however, find morphologic support for FFC formation as a resultof magmatic intrusion and sill formation. The lines of evidencefavoring magmatic intrusion as the formation mechanism includedthe large amount of floor shallowing, the unaltered crater rim crestheight, the lack of crater symmetry, moat features, the location ofFFCs away from basin edges where no basin-related thermal anom-alies are expected, and the significant population of FFCs withdiameters less than 30 km, a diameter range not favored for vis-cous relaxation (see Jozwiak et al., 2012).

Significant increases in our knowledge of FFCs from LOLA dataare high-resolution topography and detailed profiles of the craterfloor, permitting morphometric analysis of crater depth, shapeand volume, and analysis of floor-fracture structure. In contrastto fresh craters with concave-down floor profiles (such as Tycho;Fig. 1a), FFCs predominately display either flat or slightly upbowedfloor profiles (such as displayed by Humboldt; Fig. 1b), or convex-up floor profiles (as displayed by the smaller, Bohnenberger crater;Fig. 1c). Models of viscous relaxation predict that this process willresult in a shallower, gently upbowed floor (Hall et al., 1981;Dombard and Gillis, 2001; Fig. 1d), inconsistent with both theobserved broadly flat to slightly upbowed floor profiles (Fig. 1b)and the strongly upbowed floor profiles (Fig. 1c). On the otherhand, these observed profiles (Fig. 1b and c) might be explainedas different manifestations of a magmatic intrusion process, arisingfrom different laccolith morphologies, a hypothesis that we inves-tigate in Section 3.

Based on the morphologic and morphometric support for mag-matic intrusion as the formation mechanism for FFCs (e.g. Schultz,

-4.5

-4

-3.5

-3

-2.5

-2

-1.5

-1

-0.5

0

0.5

1

0 50 100 150 200 250Distance [km]

Ele

vati

on [

km]

HumboldtDiameter=207 km

Fresh Depth = 5.2 kmHumboldt Depth = 3.5 km

1.0

0.5

0.0

-0.5

-1.0

-1.5

-2.00 0.5 1.0 1.5 2.0

r / R

Hei

ght,

km

Lalande

t / = 0

t / = 0.5 x 10-8

t / = 1.5 x 10-8t / = 2.0 x 10-8

t / = 5.0 x 10-8

(fresh crater)

d

ho (43.3�S, 11.3�W), with a concave-down floor profile, (b) the large floor-fracturedthe small floor-fractured crater Bohnenberger (16.2�S, 40.0�E), with a highly domeda model of a viscously relaxed crater from Hall et al. (1981). The topographic profilesfor (a, b, and c) were measured from LOLA topography, and the fresh depths were

ater floor illustrates the exceedingly high thermal gradients (interpreted from thelar to profiles observed in craters Humboldt and Bohnenberger.

426 L.M. Jozwiak et al. / Icarus 248 (2015) 424–447

1976; Jozwiak et al., 2012), we investigate the mechanics of mag-matic intrusion to further develop and test this hypothesis. Weexplore the consequences of the intrusion process, including frac-turing associated with crater floor uplift, the cooling and degassingof the body, and the associated volcanic and solidification pro-cesses. Additionally, we evaluate the potential contribution of largesubsurface magmatic bodies to the Bouguer gravity anomalies ofFFCs.

0

1

2

3

4

5

6

0 50 100 150 200 250

Crater Diameter [km]

Cra

ter

Dep

th [

km] = Floor-Fractured Craters

= Fresh Craters

Fig. 2. Depth to diameter relationship for lunar floor-fractured craters (FFCs) (redsquares) compared with the idealized depth to diameter ratio for unmodifiedcraters of identical diameter (blue diamonds and blue line), calculated using theequations of Pike (1980) for simple craters (D < 15 km) and complex craters(D > 15 km). The population of FFCs shows significantly shallower floors than wouldbe predicted given their equivalent fresh-crater diameter. (For interpretation of thereferences to color in this figure legend, the reader is referred to the web version ofthis article.)

Table 1Characteristics of floor-fractured craters (from Schultz, 1976; Jozwiak et al., 2012).

Floor-fractured craterclass

Class characteristics

1 Radial and/or concentric fractures, crescentcrater Humboldt

2 Well-defined wall scarp, uplifted central reg

3 Wide moat between base of the crater wallGassendi

4a V-shaped moat, radial and/or concentric fra

4b V-shaped moat with pronounced inner ridgecrater Gaudibert

4c V-shaped moat, hummocky interior. Examp

5 Degraded crater walls, radial and concentric

6 Mare-flooded interiors, concentric fracture p

2. Characteristics of floor-fractured craters

2.1. Morphology

The most defining morphologic characteristic of FFCs beyondfloor fractures is the anomalously shallow floor (Fig. 1, compare1a to 1b and 1c). Fresh impact crater floor depth is commonlytaken to be a function of diameter, and follows the depth to diam-eter regression given by Pike (1980). Floor-fractured craters plotsystematically below this fresh crater regression (Fig. 2) (Schultz,1976; Jozwiak et al., 2012), and are thus defined as anomalouslyshallow. It has also been noted by both Schultz (1976) andJozwiak et al. (2012) that, despite shallower depth to diameterrelationships, FFCs preserve the fresh-crater rim crest elevationto diameter relationship given by Pike (1980). This suggests thatthe process that shallows floor-fractured craters is confined tothe crater interior, and specifically to the crater floor region.

Floor-fractured craters are divided into eight morphologic sub-classes (Schultz, 1976; Jozwiak et al., 2012) based on their specificmorphologic characteristics. Table 1 provides a brief description ofmajor FFC characteristic morphologies and Figs. 3–8 show charac-teristic craters for each morphologic class; we direct readers seek-ing a more detailed characterization to Jozwiak et al. (2012).

2.2. Spatial distribution and associations

We now examine the lunar population of FFCs as a whole, inparticular, their spatial distribution and relation to impact basins,lunar mare deposits, and crustal thickness. Fig. 9A shows the global

shaped patches of mare material adjacent to the crater wall region. Example,

ion/convex-up floor profile, concentric fractures. Example, crater Vitello

and crater interior, radial and concentric (polygonal) fractures. Example, crater

ctures, convex-up floor profile. Example, crater Bohnenberger

on the interior side, subtle fractures, irregular convex-up floor profile. Example,

le, crater at (52.7�S, 151.8�W)

(polygonal) fractures. Example, crater Von Braun

attern near crater wall. Example, crater Oppenheimer

Crater Rim Crest

Edge of Crater Wall

Mare Deposit

Wide Fracture/ Graben

Fractures

Central Peaks

BA

A

A

C

-5.0

-4.0

-3.0

-2.0

-1.0

0

1.0

2.0

0 50 100 150 200 250Distance [km]

Ele

vati

on [

km]

Humboldt

D

Crater Rim Crest

Crater Rim Crest

Central Peak

Mare Patch

Mare PatchFractures

Class 1

Fig. 3. Crater Humboldt, class 1, 207 km diameter, coordinates: 27.2�S, 80.9�E. (A) Topographic contour map with 250 m contour intervals on a color-coded altimetry map(blue low, red high). (B) Geologic sketch map: the prominent FFC morphologic features in the crater Humboldt, the most prominent of which are the complex of radial andconcentric fractures and the wall-adjacent patches of mare material. (C) LOLA gridded topographic data (512 px/deg) overlain with LROC-WAC image data for Humboldtshowing the location of the topographic profile shown in (D). (D) Topographic profile of the crater Humboldt showing similar characteristics to those in (B). (Forinterpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)


distribution of all classes of lunar floor-fractured craters super-posed on a LOLA altimetry map (Jozwiak et al., 2012), and Fig. 9Bshows all classes plotted on a map of lunar mare deposit distribu-tion and the location of all craters/basins >200 km in diameter(Head et al., 2010). Fig. 9C–J shows the distribution of membersof individual crater morphologic classes (Table 1), and the distribu-tion of lunar mare deposits and craters/basins >200 km in diame-ter. Fig. 10A shows the distribution of FFCs plotted on a map ofcrustal thickness (Wieczorek et al., 2013). Crustal thickness model1 (Wieczorek et al., 2013) was used because it incorporated thehigh porosity inferred by GRAIL (12%) and the mantle density of3220 kg/m3 was most consistent with previously inferred mantledensities (Kiefer et al., 2012). However, the specific choice of modeldoes not have a large effect on our analysis. The average crustalthickness changes, but the overall shape of the distributionremains similar. Morphologic investigations indicate that FFCswith D < 40 km have domed, convex-up floors (signifying lacco-lithic intrusions), whereas FFCs with D > 40 km have flatter floor

profiles (signifying piston-like uplift of the crater floor) (Jozwiaket al., 2012). We thus subdivide the population using diameter asa proxy for floor morphology to look for differences in areal distri-bution and correlations with crustal thickness. The distribution ofFFCs plotted on a map of crustal thickness, and showing the distri-bution of lunar maria, is shown in Fig. 10B, and FFC frequency dis-tribution as a function of crustal thickness in Fig. 10C.

Several correlations are clear: (1) Nearside/farside: The majorityof FFCs (65%) occur on the lunar nearside (Fig. 9B). Those on thelunar farside are concentrated in the South Pole-Aitken and Mos-coviense basins. (2) Association with mare basalt deposits: FFCs tendto occur along the margins of mare basalt deposits and in adjacentupland areas (Fig. 9B). There is generally a low abundance of super-posed craters in the interiors of the major maria (e.g., Imbrium,Serenitatis, Oceanus Procellarum), older craters having beenflooded and buried by the extensive mare fill. Most craters super-posed on the maria formed in the latter stages of mare fill (Eratos-thenian) or later (Copernican). Thus, these craters are less likely to

A

Crater Rim Crest

Edge of Crater Wall

Fractures

Central Peaks

B

A

A’

C

-2.5

-2.0

-1.5

-1.0

-0.5

0

0.5

0 10 20 30 40 50 60

Distance [km]

Ele

vatio

n [k

m]

Vitello

D

Crater Rim CrestCrater Rim Crest

Central Peak

Fractures Fractures

Class 2

Fig. 4. Crater Vitello, class 2, 44 km diameter, coordinates: 30.4�S, 37.5�W. (A) Topographic contour map with 100 m contour intervals on a color-coded altimetry map (bluelow, red high). (B) Geologic sketch map: the most prominent features of Vitello are the numerous concentric fractures surrounding the uplifted central peak region. (C) LOLAgridded topographic data (512 px/deg) overlain with LROC-WAC image data for Vitello showing the location of the topographic profile shown in (D). (D) Topographic profileof the crater Vitello showing similar characteristics to those in (B). (For interpretation of the references to color in this figure legend, the reader is referred to the web versionof this article.)


host extrusions and intrusions (see Fig. 12) and to form floodedcraters or FFCs. (3) Association with impact basins: There is a highlevel of correlation of FFCs with impact basin interiors (e.g., SouthPole Aitken, Moscoviense) (Fig. 9B) and impact basin margins (e.g.,Crisium, Nectaris, Humorum, Imbrium, Serenitatis). (4) Correlationwith crustal thickness: For the following discussion ‘‘thin crust’’ isdefined as <15 km, ‘‘intermediate crust’’ as between 15 km and35 km, and ‘‘thick crust’’ as >35 km. The map distribution of FFCsand crustal thickness (Fig. 10A) shows a strong positive correlationof FFC occurrences and intermediate crustal thickness (along thewestern margins of Oceanus Procellarum; along the southern mar-gins of the nearside maria, Humorum, Nubium, Nectaris; along theeastern limb maria, Crisium and Smythii). It is possible that FFCscould have formed in areas of the lunar nearside maria at lowervalues of crustal thicknesses, but have been covered by theobserved later mare fill (Fig. 10B). FFC occurrences at lower crustalthicknesses are typical of the South Pole-Aitken basin, where the

distribution of lunar maria is patchy; this FFC distribution in thin-ner crust may be an example of the appearance of nearside lunarmaria in the earlier stages of mare filling, prior to complete fillingwith mare basalts. Indeed, Fig. 10B shows that the presence of con-tinuous mare fill on the nearside occurs in areas of thinned crust(compare Fig. 10A and B) and thus may be obscuring some FFCbelow the maria.

A quantitative assessment is obtained from examining the FFCfrequency distribution as a function of crustal thickness(Fig. 10C). The mean of the FFC population occurs at a crustal thick-ness of 26 km and the median at 28 km. A peak in the FFC distribu-tion occurs in regions of crustal thickness �26–34 km, which isnearly coincident with the average nearside crustal thickness of�31 km, and bounded by the average lunar crustal thicknessof �34 km and the average crustal thickness beneath the mare of�24 km. Only ten FFCs occur in crust thicker than the averagefarside crustal thickness (omitting South Pole-Aitken basin) of

Crater Rim Crest

Edge of Steep Wall

Moat Feature

Central Peaks

Floor Fractures

Wide Fractures/ Graben

-3.0

-2.5

-2.0

-1.5

-1.0

-0.5

0

0 20 40 60 80 100 120 140 160

Distance [km]

Elev

atio

n [k

m]

GassendiCrater Rim Crest

Crater Rim Crest

Moat

Central Peaks

Wall TerracesFractures

A A’

A

A’

A

C

B

D

Class 3

Fig. 5. Crater Gassendi, class 3, 110 km diameter, coordinates: 17.5�S, 39.9�W. (A) Topographic contour map with 250 m contour intervals on a color-coded altimetry map(blue low, red high). (B) Geologic sketch map: the most prominent features of Gassendi are the complex of radial and concentric fractures and the large moat feature adjacentto the southern rim of the crater. (C) LOLA gridded topographic data (512 px/deg) overlain with LROC-WAC image data for Gassendi showing the location of the topographicprofile shown in (D). (D) Topographic profile of the crater Gassendi showing similar characteristics to those in (B). (For interpretation of the references to color in this figurelegend, the reader is referred to the web version of this article.)


�40 km; these occur in the distal regions of some farside cratersand basins (e.g., north of Tsiolkovsky and outside, but ringing, Ori-entale). A small population of FFCs is found on crustal thicknesses<10 km, and these correspond to FFCs located on the floors ofimpact basins, such as South Pole-Aitken and Moscoviense, whichhave excavated deep into the lunar crust, but have not been asextensively flooded by mare lavas as the nearside maria (compareFig. 10A and B).

These correlations permit us to combine: (1) our model ofmagma ascent and eruption in relation to crustal structure andoverpressurization of ascending dikes (see Wilson et al., 2014;Head and Wilson, 1989, 1992), and (2) the distribution and charac-teristics of the population of floor-fractured craters (Jozwiak et al.,2012 and the work reported here). We first simplify the range ofdike emplacement events in terms of a population of overpressur-ization values. As described in Head and Wilson (1992) (see alsoYingst and Head, 1997), we adopt the concept that dikes propagat-ing from magma reservoirs will be characterized by a range ofoverpressurization values (Fig. 11A), reflecting variations in therates of magma arriving in the reservoir and the ability of the res-ervoir to deform elastically to accommodate the increased strain.

We conceptualize this as a normal frequency distribution(Fig. 11A) in the following manner: Starting with the assumptionthat dikes are propagating through a relatively thin crust, the high-est overpressurization values (OP 1) will be most likely to formdikes that will penetrate to the surface and form effusive eruptions(Fig. 11B). Intermediate overpressurization values (OP 2) willapproach the surface and be more likely to form intrusive sillsrather than extrusive flows. The lowest overpressurization values(OP 3), are sufficient to propagate dikes from the reservoir, butinsufficient to approach the surface; these latter types (OP 3) arethe most likely to form solidified dikes, concentrated lower inthe crust (Fig. 11B).

We now examine the relationship of this model to implicationsfor the process of ascent and emplacement of dikes and predictedvariations in crustal thickness. We first describe a series of geolog-ical and crustal thickness settings (Fig. 11C) and make predictionsabout the relation of overpressurization values and crustal thick-ness for each of these settings. We then examine the distributionand characteristics of the actual population of FFCs (Figs. 9 and10) in order to assess whether they fit the predictions of the sim-plified conceptual model. Clearly, the actual situation is likely to

Moat Feature

Floor Fractures

Wide CentralFractureCrater Rim

Crest

Edge of Steep Wall

A B

D

A A’

C-3.5

-3.0

-2.5

-2.0

-1.5

-1.0

-0.5

0

0 5 10 15 20 25 30 35 40 45

Distance [km]

Ele

vatio

n [k

m]

A A’

Bohnenberger

Crater Rim CrestCrater Rim Crest

“V” Profile Moat“V” Profile Moat

Central Fracture

Class 4a

Fig. 6. Crater Bohnenberger, class 4a, 33 km diameter, coordinates: 16.2�S, 40.0�E. (A) Topographic contour map with 100 m contour intervals on a color-coded altimetry map(blue low, red high). (B) Geologic sketch map: the most prominent features of Bohnenberger are the ‘‘v’’-profile moat at the crater wall interface, and the wide fracture acrossthe center of the crater floor. (C) LOLA gridded topographic data (512 px/deg) overlain with LROC-WAC image data for Bohnenberger showing the location of the topographicprofile shown in (D). (D) Topographic profile of the crater Bohnenberger showing similar characteristics to those in (B). (For interpretation of the references to color in thisfigure legend, the reader is referred to the web version of this article.)


be more complex than this simple model (e.g., thermal variations,mantle heterogeneity, etc.), but it is instructive to assess where thebasic model is successful in making predictions and where it fails.

Setting 1: Impact basin interiors in crust of intermediate regionalthickness (Fig. 11C): In this crustal thickness environment, lowercrustal thicknesses are predicted to favor extrusions over intru-sions for OP 1 dikes (Fig. 11A and B), and this results in a seriesof effusive lava flows topographically contained within the basins.Supporting evidence for this is the extensive lava flow unit stratig-raphy of mare basin fill (e.g., Hiesinger et al., 2011) and associatedfeatures suggesting high effusion rates (e.g., sinuous rilles; Wilsonand Head, 1981; Hurwitz et al., 2013). OP 2 events could produceintrusions and FFCs in the basin interiors in the early stages offlooding, but are likely to be buried by subsequently accumulatinglava.

Examples of these types of settings and occurrences areobserved in Mare Imbrium, Serenitatis, Crisium, and OceanusProcellarum (Fig. 10A and B).

Setting 2: Impact basin rims (Fig. 11C): In these cases, magmarising in dikes is predicted to encounter the locally thickened crust

formed in association with impact basin rims. OP 1 dikes risingunder craters superposed on the basin rims are favored to formsills, rather than extrusions. Loading of the lithosphere by the adja-cent volcanic load is also predicted to create a stress state and dis-tribution that favors rim locations (Solomon and Head, 1980).

Examples of these types of occurrences (Fig. 10A and B) includeGassendi on the rim of the Humorum basin, Posidonius on the mar-gin of the Serenitatis basin, Cleomedes on the edge of the Crisiumbasin and Schluter on the rim of the Orientale Basin.

Setting 3: Crust of intermediate thickness (Fig. 11C): In these crus-tal thickness environments, OP 1 (Fig. 11A and B) dikes are pre-dicted to typically form intrusions, rather than extrusions, andthus this is the generally preferred environment for FFC develop-ment. OP 2 dikes will typically have insufficient overpressurizationto form sills and will solidify as dikes. As crust becomes thicker,typical OP 1 dikes are predicted to have insufficient overpressur-ization (and thus intruded volume) to create major thick sills inlarge craters. Although OP 1 dikes could form thin sills beneathlarge craters in this setting, there is likely to be insufficient magmaoverpressure and volume to create the vertical inflation necessary

A

Crater Rim Crest

Edge of Steep Wall

Fractures

Wide Fracture/Graben

B

C

A A’

-3.0

-2.5

-2.0

-1.5

-1.0

-0.5

0

0 10 20 30 40 50 60 70

Distance [km]

Ele

vatio

n [k

m]

Von Braun

A ACrater Rim Crest

Crater Rim Crest

FracturesWide Fracture

D

Class 5

Fig. 7. Crater Von Braun, class 5, 60 km diameter, coordinates: 41.1�N, 78.0�W. (A) Topographic contour map with 100 m contour intervals on a color-coded altimetry map(blue low, red high). (B) Geologic sketch map: the most prominent features of Von Braun are the numerous radial fractures in the center of the crater floor and the wall-adjacent concentric fractures. (C) LOLA gridded topographic data (512 px/deg) overlain with LROC-WAC image data for Von Braun showing the location of the topographicprofile shown in (D). (D) Topographic profile of the crater Von Braun showing similar characteristics to those in (B). (For interpretation of the references to color in this figurelegend, the reader is referred to the web version of this article.)


to form thick sills, and thus the deformation necessary to producefloor-fractured craters. OP 1 dikes intruded into smaller craters inthis crustal thickness environment, however, are predicted to havesufficient overpressurization (and volume) to upbow crater floors.Thus, smaller diameter FFCs are predicted to be favored to form inthis crustal environment, compared with larger FFCs.

Examples of these types of occurrences (Fig. 10A and B) aboundalong the western margins of Oceanus Procellarum where numer-ous FFCs occur in a broad swath of crust of intermediate thickness.Furthermore, as predicted by theory (Fig. 11C), in this region FFCswith diameters >40 km occur preferentially toward the regions ofthinner crust (the maria); smaller FFCs (diameters <40 km) occurpreferentially toward the thicker crust (farside highlands) (seeFig. 10A and B and compare with Fig. 11A and B). Away from theedges of the maria (southern highlands) zones of intermediate

crust are characterized by much less abundant large and small FFCs(Fig. 10B), probably related to the intrusion of a few OP 1 dikes.

Setting 4: Thick crust (Fig. 11C): In these crustal thickness envi-ronments, OP 1 dikes (Fig. 11A and B) are predicted to stall atdepths sufficiently below the surface that they do not intersectthe low density breccia zones associated with impact craterssuperposed on the crust. OP 2 and OP 3 dikes stall and solidify evendeeper in the crust.

Examples of these types of environments and settings occur onthe lunar farside, which is characterized by typically higher crustalthicknesses (Fig. 10A and B); these regions are notably deficient inFFCs (Fig. 10A).

Setting 5: Peak-ring basins in thick crust (Fig. 11C): These basinscreate a locally thinned crustal environment in an otherwise thickcrust. In this setting, OP 1 dike events are now predicted to be

A

Crater Rim Crest

Edge of Steep Wall

Fractures

CraterOppenheimer H

CraterOppenheimer U

Dark Halo Craterand Deposit

Dark MantlePyroclastic Deposit

Vent

B

A

A’

C

-6.0

-5.0

-4.0

-3.0

-2.0

-1.0

0

0 50 100 150 200 250 300

Distance [km]

Ele

vatio

n [k

m]

Oppenheimer

A A’

Crater Rim Crest

Crater Rim Crest

Crater Rim Crest

Crater Rim Crest

Oppenheimer H

Oppenheimer U

D

Class 6

Fig. 8. Crater Oppenheimer, class 6, 208 km diameter, coordinates: 35.2�S, 166.3�W. (A) Topographic contour map with 250 m contour intervals on a color-coded altimetrymap (blue low, red high). (B) Geologic sketch map: the most prominent features of Oppenheimer are the numerous wall-adjacent concentric fractures. Additionally the darkhalo crater and deposit in the southern part of the crater floor, and the vent and dark mantle pyroclastic deposit in the northern part of the crater are important morphologicfeatures. Oppenheimer also hosts two superposed craters, Oppenheimer H and U, which themselves display morphologic features of floor-fractured craters and extensivepyroclastic deposits (Gaddis et al., 2013). (C) LOLA gridded topographic data (512 px/deg) overlain with LROC-WAC image data for Oppenheimer showing the location of thetopographic profile shown in (D). (D) Topographic profile of the crater Oppenheimer showing similar characteristics to those in (B). (For interpretation of the references tocolor in this figure legend, the reader is referred to the web version of this article.)


capable of reaching superposed impact crater low-density brecciazones, and floor fractured craters are produced. Depending onthe crustal thickness and exact overpressurization values, a floodedcrater, rather than a FFC, might result. In some cases, both might beobserved.

Examples of these types of FFC occurrences are common in theMoscoviense basin on the lunar farside (Fig. 10A). Here, thinnercrust is correlated with Mare Moscoviense and the FFC Komarov,and several smaller examples within the maria. In contrast to mostnearside mare-flooded impact basins, Mare Moscoviense floodsonly the inner depression and the northeast portion of the adjacentplatform (Kramer et al., 2008; Thaisen et al., 2011), potentiallyrevealing an example of the earlier stages of flooding of many near-side basins. Komarov, a large FFC (�78 km) occurs on the basin

interior platform above the floor of Mare Moscoviense (seeFig. 11C, setting 2), whereas the smaller FFCs occur in the mareinterior. Several small unnamed FFCs occurring in peak-ring basinsin thicker crust include the farside basins Freundlich-Sharonov andCoulomb-Satron (e.g., Baker et al., 2011, 2012; Jozwiak et al., 2012).

Setting 6: Large multi-ring basins in regionally thick crust: In thiscase, the crust is anomalously thin relative to the setting of theadjacent regional thick crust, and a variety of magmatic and volca-nic features are predicted in the basin interior. As in Setting 1(impact basin interiors), OP 1 dikes are predicted to propagate tothe surface and to form extrusive lava flows, whereas OP 2 dikesare predicted to produce sills and FFCs.

The principle example of this type of occurrence (the SPA-SouthPole Aitken basin) is located in the regionally thick crust of the

135° W

135° W 135° E

135° E90° W

90° W

90° E

90° E

90° N 90° N

90° S90° S

45° W

45° W

45° E

45° E

45° N 45° N

45° S 45° S

0°

0°

0° 0°

180° W

180° W

180° E

180° E

Class 1 Class 2 Class 3 Class 4a

Class 4b Class 4c Class 5 Class 6

A

Class 1 Class 2 Class 3 Class 4a

Class 4b Class 4c Class 5 Class 6

135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

45°N

0°

45°S

90°S180° W

180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E90°N

45°N

0°

45°S

90°SB

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°NClass 1

Oceanus Procellarum

Orientale

South Pole-Aitken South Pole-Aitken

Nectaris

ImbriumSerenetatis Crisium

Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

C

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°NClass 2

Oceanus Procellarum

Orientale

South Pole-AitkenSouth Pole-Aitken

Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

D

90°N

Fig. 9. Global distribution map of the lunar population of FFCs as listed by Jozwiak et al. (2012). The total number of craters mapped is 170. (A and B) All FFCs. (C) Class 1 FFCs,e.g. Humboldt. (D) Class 2 FFCs, e.g. Vitello. (E) Class 3 FFCs, e.g. Gassendi. (F) Class 4a FFCs, e.g. Bohnenberger. (G) Class 4b FFCs, e.g. Gaudibert. (H) Class 4c FFCs, e.g. crater at(52.7�S, 151.8�W). (I) Class 5 FFCs, e.g. Von Braun. (J) Class 6 FFCs, e.g. Oppenheimer. (A) Color-coded altimetry map (blue low, red high), LOLA gridded 512 px/deg. (B–J) USGSgeologic map of mare units and outline of all lunar craters with diameter >200 km (Head et al., 2010). (For interpretation of the references to color in this figure legend, thereader is referred to the web version of this article.)


0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°NClass 3

Oceanus Procellarum

Orientale


Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

E

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°NClass 4a

Oceanus Procellarum

Orientale


Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

F

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°NClass 4b

Oceanus Procellarum

Orientale


Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

G

Fig. 9 (continued)


lunar farside (Fig. 10A and B). Because of the ancient age of the SPA,a significant number of superposed craters are observed on thefloor of the basin in the size range typical of FFCs. Impact basinssuperposed on the SPA basin interior (e.g., Apollo, Schrodinger) fur-ther thin the crust and are predicted to result in lava flooding andintrusion, which is indeed observed. In the crustal thickness typicalof the SPA basin interior, OP 1 dike events are predicted to result inextrusion of lavas and filling of crater interiors, as amply seen inIngenii, Van de Graff, Von Karmen, Jules Verne, Leibnitz, Apollo,and in adjacent intercrater areas (Yingst and Head, 1997, 1999).OP 2 dike events are predicted to result in shallow intrusion andsill formation, creating FFCs. Indeed, a wide variety of floor-frac-tured craters are observed across the basin floor (Oppenheimer,Fizeau, Haret) and in the interior of the superposed Apollo basin,consistent with these predictions (see also, Yingst and Head,

1997). Notable in the SPA FFC population is that examples of smal-ler FFCs (<40 km diameter) are more common than larger FFCs.Lower overall pressurization values would tend to favor lower vol-ume intrusions and thus smaller FFCs, a factor that might be con-sistent with the minimal amount of general flooding.

A remaining question is: Why does the unusually thin crust inthe SPA basin (Fig. 10C) result in smaller amounts of mare floodingand fill, and more FFCs, than in other nearside basins, such asImbrium, Serenitatis, and Crisium? Several possible explanationsexist: (1) the thermal structure on the nearside and farside coulddiffer, resulting in thermal barriers to magma reservoir formation,as well as density barriers; this, in turn, could produce variations inmagma ascent and emplacement (Head and Wilson, 1992; (2) thebroad context of the thicker crust in which the SPA basin occurscould influence magma ascent and eruption; (3) the residue of

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

Class 4c

Oceanus Procellarum

Orientale


Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

H

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°NClass 6

Oceanus Procellarum

Orientale


Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

J

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

Class 5

Oceanus Procellarum

Orientale

South Pole-Aitken South Pole-Aitken

Nectaris


Smythii

Australe

Moscoviense

Humorum

Humboldtianum

Freundlich-Sharonov

I

Fig. 9 (continued)


the huge SPA basin-forming event (the impact melt sea, Vaughanet al., 2013; and the differentiation of this sea, Vaughan andHead, 2014) could produce a distinctive solidified melt-sheet den-sity and porosity structure that would influence the ascent anderuption of magma. Changes in the density structure of the crust,as a result of the differentiation of an impact melt sea coulddiscourage shallow melt stalling. Changes in porosity, specificallya porosity decrease resulting from the crystallization of the impactmelt sea into a competent layer could also inhibit dike propaga-tion. The presence of large fractures decreases the overall layerstrength, thus dikes would require a higher overpressurization topropagate through the crystallized melt sheet layer; (4) thestructure and composition of the sub-melt sheet mantle residue

(fractures, injected melt veins, partial melting products, etc.), couldhave further influenced the nature and density structure of thesubstrate in this region; (5) differences in the much deeper marebasalt magma generation areas could be playing a role in magmareservoir sizes, overpressurization values, and ascent and eruption.We are currently assessing the influence of solidified impact meltseas (Vaughan et al., 2013; Vaughan and Head, 2014) on theestimates of crustal thickness used in this study (Wieczoreket al., 2013), and on the subsequent ascent, intrusion and eruptionof magma in the SPA basin.

Setting 7: Ancient viscously relaxed basins: In some cases, impactbasins will have lost a significant amount of their topographicexpression and crustal thickness variation due to viscous

Diameter > 40 kmDiameter < 40 km

0° 45°E 90°E 135°E 180°E45°W90°W135°W

0° 45°E 90°E 135°E 180°E45°W90°W135°W

90°N

45°N

0°

45°S

90°S

90°N

45°N

0°

45°S

90°S180°W

180°W

Crustal ThicknessHigh: 67 kmLow: 1 km

A

Diameter > 40 kmDiameter < 40 km

0° 45°E 90°E 135°E 180°E45°W90°W135°W

0° 45°E 90°E 135°E 180°E45°W90°W135°W

90°N

45°N

0°

45°S

90°S

90°N

45°N

0°

45°S

90°S180°W

180°W

Crustal ThicknessHigh: 67 kmLow: 1 km

B

0

5

10

15

20

25

30

35

40

45

50

5 10 15 20 25 30 35 40 45 50 55 60 65 70Crustal Thickness [km]

Freq

uenc

y

Average Global Crustal Thickness

Average Farside Crustal Thickness (omitting SPA)

0

Basin floors(e.g. SPA,Smythii, Moscoviense) Adjacent to farside basins

(e.g. Tsiolkovsky crater)

FFC Crustal Mean

FFC Crustal Median Average Nearside Crustal Thickness

Average Crustal ThicknessBeneath Maria

C

Fig. 10. Global distribution of lunar FFCs (Jozwiak et al., 2012) and frequency distribution of associated crustal thicknesses. (A and B) Global distribution of FFCs divided bycrater diameter, D > 40 km black circles and D < 40 km white squares. (C) frequency distribution of crustal thickness beneath FFCs. Average lunar crustal thickness 34 km(black dashed line), average nearside crustal thickness 30 km (red dashed line), average farside crustal thickness, omitting South Pole-Aitken basin, 40 km (blue dashed line),mean FFC crustal thickness 26 km (green dashed line), median FFC crustal thickness 28 km (purple dashed line). FFCs with a crustal thickness below 15 km are located withinlarge basins such as South Pole-Aitken and Moscoviense. FFCs with a crustal thickness greater than 40 km are located near, but outside large farside craters and basins such asTsiolkovsky. (A and B) Color-coded crustal thickness basemap derived from GRAIL data (blue low, red high), Wieczorek thickness model 1 (Wieczorek et al., 2013), minimumcrustal thickness set a priori to 1 km. (B) USGS geologic map of mare units. (C) Crustal thickness bin width of 5 km. (For interpretation of the references to color in this figurelegend, the reader is referred to the web version of this article.)


relaxation; such basins are typically old and thus will also haveaccumulated a large number of superposed craters. In this interme-diate crustal thickness environment, OP 1 dikes are predicted toform intrusions and extrusions, and OP 2 dikes will produceintrusions or simply solidify in the crust as dikes.

Mare Australe, filling the ancient Australe basin, is one exampleof this situation (Fig. 10A and B). Rather than extensive and

continuous lava plains filling the basin interior, such as is observedin the younger Imbrium, Serenitatis and Crisium basins, MareAustrale is characterized by mare filling of a large number of indi-vidual craters superposed on the basin interior. The lack of FFCs inAustrale, and the abundance of individual craters flooded withmaria, support the interpretation that OP 1 dikes are intruding intothe crater interiors and forming eruptions rather than stalling to

Dike FrequencyDistribution

Dik

e O

verp

ress

uriz

atio

n

A

Lunar Surface

OP 3 OP 3

OP 2 OP 2

OP 1

B

OP 3Solidified

Dike

OP 2Shallow Intrusion,

Sill, Floor-FracturedCrater

OP 1Accumulated

Extrusive Flows

OP 3OP 1OP 2

C2 2222

1 11111 1 111

3 3 333

Setting 2Impact Basin

Rims

Setting 1Impact Basin

Interiors

Setting 3Crust of

IntermediateThickness

Setting 4Thick Crust

Setting 5Peak-Ring Basins

In Thick Crust

Fig. 11. (A) Schematic of the frequency distribution of dikes propagating toward the surface resulting from overpressurization of the magma reservoir. Horizontal linerepresents the lunar surface. The most overpressurized dikes (OP 1) will reach the surface to form flows, while the least pressurized (OP 3) will solidify deeper in the crust; OP2 dikes may form shallow intrusions such as FFCs. Despite having the highest mean magma overpressurization, OP1 dikes likely represent only a few percent of the totalnumber of intruded dikes. (B) Cross-sectional representation of the consequence of the three dike overpressurization types. (C) Crustal cross-sectional representation ofdifferent crustal thickness and feature/structure settings in relation to the global distribution of floor-fractured craters and crustal thickness (Figs. 9 and 10). See text fordetailed discussion.


form sills. Some of these craters may have begun as floor-fracturedcraters, but exceeded the sill-forming overpressure sufficiently toextrude out and flood the crater floor. Such a setting (e.g., initialintrusion, sill formation, floor uplift, floor fracturing, and craterinterior flooding) suggests very finely tuned overpressurizationconditions.

In summary, the general paradigm of magmatic overpressuriza-tion levels and implications for (1) eruption, (2) sill intrusion andfloor-fractured crater formation and (3) dike cooling and solidifica-tion (Fig. 11A and B), when correlated with crustal thickness pre-dictions (Fig. 11C) and observations (Fig. 10A–C), provides aplausible interpretation for the vast majority of floor-fractured cra-ter occurrences.

2.3. Temporal distribution

It is difficult to ascertain the age of the deformation process thatproduced FFCs because (1) most of the activity occurs in the sub-surface, (2) fracturing is confined to the crater interiors, and (3)deposits (small mare patches and dark-halo craters) are too smallto obtain reliable impact crater size-frequency distribution ages.By analyzing the ages of the host craters, however, we can deter-mine if the process operated solely in the early part of lunar his-tory, or if it spanned the entire cratering record. The crater ageperiods and individual crater ages were taken from Losiak et al.(2009). The frequency distribution of crater formation age for allcraters characterized by fractured floors (FFCs) with D > 30 km is

0° 45° E 90° E 135° E 180° E45° W90° W135° W180° W90°N

45°N

0°

45°S

90°S180° W 135° W 90° W 45° W 0° 45° E 90° E 135° E 180° E

90°S

45°S

0°

45°N

90°N

B

Pre-Nectarian Nectarian Lower Imbrian Upper Imbrian Eratosthenian Copernican

Num

ber

of F

loor

-Fra

ctur

ed C

rate

rs

Age of Crater

Age Span of Mare Volcanism

0

35

30

25

20

15

10

5

A

Fig. 12. (A) Frequency and (B) spatial distribution of crater formation stratigraphic ages for all FFCs with D > 30 km (Losiak et al., 2009). (B) USGS geologic map of mare unitsand outline of all lunar craters with diameter >200 km (Head et al., 2010).


shown in Fig. 12A. It is clear that floor-fracturing occurs in cratersof a wide range of ages. One crater (Taruntius) was originallymapped as Copernican in age on the basis of its bright ray system.However, Hawke et al. (2004, 2006) have shown the Taruntius raysto be of compositional origin and not related to maturity, whichwas the basis for assignment to the Copernican System; they thusreassigned Taruntius to the Eratosthenian. These relationships(Fig. 12) imply that FFC formation processes are not restricted toa specific time period in early lunar history (such as the Nectarian),because craters that formed in the Lower Imbrian, Upper Imbrian,and Eratosthenian are also floor-fractured. On the other hand, thefloor fracturing process could readily operate over a more limitedtime frame than the range of different-aged craters it deformed;for example, if floor-fracturing is due to intrusion by basaltic mag-mas, then the range of formation ages may be much more closelyapproximated by the span of mare volcanism (e.g., predominantlyLower Imbrian to late Eratosthenian; Hiesinger and Head, 2006;Hiesinger et al., 2012) than by the range of ages of the craters thatthe magmas intruded. This latter interpretation, the correlation ofintrusion and floor-fracturing with the period of mare volcanism,rather than crater age (Fig. 12A), is supported by close geographiccorrelation of the FFC population with the lunar maria (Fig. 9) andthe relatively random distribution of FFCs of different ages(Fig. 12B).

The lack of a widespread, more random distribution of floor-fractured craters (Fig. 9) clearly argues against hypotheses pro-posed very early in the study of the Moon, that the formation offloor fracturing was related to the impact event itself, or to a rela-tively short term response of the crust and lithosphere to the event.

Despite the presence of hundreds of examples of ancient pre-Imbrian impact craters (Head et al., 2010), only a few are charac-terized by floor fracturing even though the thermal structure dur-ing this early period should have been more conducive to sucheffects. Indeed, the only current studies which propose impact-induced melting, do so for basin-scale cratering events (e.g.Elkins-Tanton et al., 2004). Additionally, observations of pre-Nec-tarian and Nectarian aged craters show that the fractures do nothave the same degradation state as the rest of the crater morphol-ogy, suggesting different formation times.

To summarize the major characteristics of FFCs: (1) FFCs pos-sess shallow floors, as represented by their crater depth to diame-ter ratio (Fig. 2), (2) the process which formed FFCs does not extendpast the crater floor, as demonstrated by the unaltered crater rimcrest to diameter ratio (Schultz, 1976; Jozwiak et al., 2012), (3)FFCs contain a variety of morphologic features including variouspatterns of floor fractures, patches of mare material, vents and darkhalo deposits, and both domed (convex-up) and flat floors (Figs. 3–10). As was discussed by Jozwiak et al. (2012) these morphologiesfavor FFC formation via magmatic intrusion and sill formation, aprocess that we now investigate in more detail.

3. Magmatic intrusion and sill formation beneath an impactcrater

We propose (and subsequently test) a magma-based hypothesisfor the formation of FFCs, as follows. The hypothesized process ofmagmatic intrusion and sill formation (Fig. 13) involves the

A

B

C

D

E

Propagating Dike, ρm

Lunar Crust, ρcBrecciated Region, ρb

Faults

Fig. 13. Schematic cross-section illustration the process of FFC formation via magmatic intrusion and sill/laccolith formation. (A) Dike propagates towards the surface withenough driving pressure to propagate to the surface. (B) Vertically propagating dike encounters highly fractured, low density region beneath crater, stalls. (C) Driving pressureof the now-stalled dike exceeds the local lithostatic pressure; dike begins to propagate laterally, forming a sill. The sill extends to edges of crater floor, where increasedoverburden pressure of the crater wall and crater rim cause a pinching off of the dike, and cessation of lateral propagation. (D) Sill continues to fill with magma. The ratio ofintrusion diameter to intrusion depth is smaller than a critical value, resulting in an overall domed, laccolith shape, and increased bending stresses over the uplifted regions ofthe crater floor. (E) The ratio of the intrusion diameter to the intrusion depth exceeds the critical value, and the sill is able to tabularly deform the overlying crater surface,localizing bending stresses at the periphery of the intrusion, and resulting in fracturing in this wall-adjacent region and piston-like uplift of the entire crater floor.


propagation of a dike from a mantle source region to the surface(Wilson and Head, 1981; Head and Wilson, 1992), and intersectionwith a crater substructure; the dike then stalls beneath the craterand spreads laterally following the dimensions of the crater floor,resulting in a sill. The sill then inflates forming a laccolith, withthe laccolith bowing up the overlying crater floor, or in some casescausing an overall piston-like uplift of the overlying crater floor(Schultz, 1976). This process is shown schematically in Fig. 13,and we note that the process of sill formation on the Moon issimilar in principle to sill intrusion on Earth (Wilson and Head,1981; Head and Wilson, 1992; Wichman and Schultz, 1996).Throughout the process, the magma movement is governed bypressure constraints—lithostatic pressure and excess pressure(pressure in the magma reservoir in excess of lithostatic), thedetailed values of which vary depending on the specific subcraterenvironment. The outcome of our modeled intrusion processesallows for variations in these parameters, without having tospecifically prescribe the exact geometry and detailed densitystructure of the subcrater environment, which is poorly known(e.g. Dence et al., 1977).

3.1. Initial dike propagation

The initial stage of this process involves the propagation of adike from a magma source at the top of the mantle through thelunar crust (Wilson and Head, 1981), as depicted in Fig. 13A. Inorder for this process to occur, the stress intensity at the upper dike

tip must exceed the local fracture toughness, and the absolutepressure in the magma source at the base of the dike must be greatenough to support the weight of the negatively buoyant magma tothe required intrusion level (Head and Wilson, 1992). The impor-tant parameters governing this process are the height of the diketip above the mantle source region, a; the density difference of hostrock minus intruding magma, Dq; acceleration due to gravity, g;and the fracture toughness of the material into which the dike isintruding, Kc. The value of fracture toughness used in this analysisis based on the field-derived values of Rubin and Pollard (1987)and is essentially identical to the value found by Parfitt (1991).The densities for lunar basaltic magma, lunar crustal material,and the lunar breccia lens material are taken from the Kieferet al. (2012) analysis of Apollo lunar samples. The driving pressure,Pd, necessary to propagate the dike to height a, modified fromRubin and Pollard (1987) is given by:

Pd ¼Kc � ððp�1 þ 0:25Þ � g � DqÞa3=2

a1=2 : ð1Þ

Equating a to the lunar crustal thickness yields the magma drivingpressure necessary for a dike to propagate from the uppermostlunar mantle to the surface. Using the material properties listed inTable 2 and 20.6 MPa is required for a dike to reach the surface,based on a �34 km mean crustal thickness (Wieczorek et al., 2013).

It is also possible that the dike in question does not have suffi-cient driving pressure to reach the surface. Head and Wilson(1992) have shown that the maximum height to which magma


in a dike can rise is equal to the hydrostatic equilibrium height ofthe dike plus the additional height produced by the driving pres-sure of the magma. To continue the example scenario from above,the hydrostatic equilibrium height of the dike is 27.4 km and theadditional height from the driving pressure is 3.9 km for a totalheight of 31.3 km, well below the presumed crustal thickness of34 km. It is worth noting that this final stalling depth is dependenton the input densities; for example, lower density magmas willstall at shallower depths and a decrease in crustal density (dueto impact-related brecciation, for example) will result in stallingat greater depths.

3.2. Dike tip and breccia lens interactions

Illustrated in Fig. 13B is the process by which the dike tipencounters the underdense breccia region beneath the crater floor,decelerates at a specific depth influenced by the density differ-ences of the brecciated region, and subsequently halts. To beginexamining the viability of the breccia lens acting as a density bar-rier preventing dike propagation, we first note that the exact nat-ure of the crust/breccia lens transition is unknown, but isunlikely to be a single specific contact. Rather, this is a compactregion marked by the density change from unaltered crustal mate-rial to heavily fractured impact breccia material related to the hostcrater, first noted in structural studies of terrestrial impact craters(e.g. Moore, 1976; Dence et al., 1977; Pohl et al., 1977). As such, thedike tip will propagate for some distance into the breccia regionuntil rise is no longer energetically favorable, that is, the Pd ofthe dike is insufficient to continue propagation through brecciamaterial. In the preceding section we calculated that a dike intrud-ing into a 34 km crustal column of uniform density will stall at�31 km for the given parameters. If we alter the crustal columnto a 24 km column of average density crust overlain by a 10 kmlayer of less dense heavily brecciated material and assume thesame driving pressure, the new stalling depth of the dike is�30 km. Thus, all other factors being equal, the smaller densityof the breccia lens means that dike propagation will be halted ata greater depth compared with intrusion into lunar crust far awayfrom a fresh crater breccia lens.

3.3. Sill formation

When vertical propagation of the dike is halted, there may besignificant driving pressure at the tip of the dike. If the drivingpressure at the dike tip exceeds the local lithostatic pressure, hor-izontal propagation will occur, forming a sill approximately circu-lar in plan-view (Rubin and Pollard, 1987). Continuing with theexample of the crater Gassendi, the lithostatic pressure at thecrust/breccia transition is approximately 16 MPa; this is exceededby the driving pressure of the dike required to reach the base of thebreccia zone, permitting fracturing in this plane and horizontalpropagation of magma as a sill.

The sill then propagates to the edge of the crater floor regionbefore it pinches off as a direct response to changes in lithostaticoverburden pressure. While the sill is beneath the breccia lens, itexperiences a relatively small overburden pressure stemming fromthe less dense breccia material and the deficit of crustal materialthat had been excavated by the impact event and is now missingfrom the crater interior (Fig. 13C). The overburden pressure atthe edge of the breccia lens/crater floor region is significantlyincreased by the presence of the crater wall and rim, a crustal col-umn consisting of somewhat denser crustal rock, and the fullextent of the regional crustal material consisting of uplifted craterrim material and superposed ejecta. For a crater with a depth of2 km, the magnitude of this overburden pressure increase in agiven crustal column is �8 MPa; this increased overburden

pressure combined with the lithostatic pressure exceeds the initialsill driving pressure, and causes the sill to pinch off at the edge ofthe crater floor. Supporting this interpretation, morphologic datafrom floor-fractured craters show no evidence for sill-related upliftextending past the crater floor (Figs. 3–8). Thus we infer that thesill propagates laterally along the base of the breccia lens formingan intrusion that is approximately circular in plan view, and mir-rors the extent of the crater floor (Fig. 13C). Complementary toour geologic observations, Thorey and Michaut (2014) have pro-duced models of magmatic intrusion beneath a crater by modelingthe deformation produced by deforming an elastic sheet pinned atthe edges. This idealized model includes the assumptions of purelyelastic deformation of the crater floor and simplifications to theinitial crater and subcrater morphology to arrive at its results.The results reproduced the two end-member morphologies of abell-shaped laccolith, and a thicker slab-like intrusion, which arein agreement with our morphologic observations of the lunar FFCpopulation using floor morphology to investigate intrusion mor-phology. Additionally all of the models show that the intrusion willnot spread laterally beyond the crater floor (Thorey and Michaut,2014) due to an increase in lithostatic pressure beneath the craterwall zone. This is in agreement with the morphologic observationthat the rim crest heights vs. diameter relationships for FFCs areindistinguishable from normal complex craters (Schultz, 1976),suggesting that the intrusion dimensions do not extend beyondthe crater wall.

3.4. Sill to laccolith transition

We now consider the transition from sill to laccolith and exam-ine the end-member intrusion morphologies. The idealized defor-mation of strata overlying an intrusion can take two forms: adomed laccolith with thin tapered edges (Fig. 13D) or a thick sillwith blunt, snub-nosed edges (Fig. 13E) (Corry, 1988). Themechanics of this process are initially described in Johnson andPollard (1973), with theoretical modeling of laccolith formationshown in Michaut (2011), and a likewise-theoretical model of lac-colith formation beneath a crater shown in Thorey and Michaut(2014). Although these two models (Fig. 13D and E) both describethe deformation of a uniform elastic sheet, and are oversimplifiedidealizations of the subcrater environment, we will consider themas possible end-members in the evolution of FFC-forming mag-matic intrusions. A critical parameter in determining the transitionbetween a domed laccolith and a sill is the scaling relationshipbetween the intrusion radius and the thickness of the overburden(Jackson and Pollard, 1988). When the radius of the intrusionreaches a critical distance relative to the overburden thickness,stresses will concentrate at the periphery, allowing fracturingalong the edges, and piston-like uplift of the overburden, resultingin a flat, sill-like intrusion with snub-nosed edges (Johnson andPollard, 1973; Rubin and Pollard, 1987). Conversely, when theradius of the intrusion is confined to a small lateral extent, theresult is thickening of the central part of the intruded region, anda pronounced doming of the overlying strata (Johnson andPollard, 1973, 1988). Jackson and Pollard (1988) approximate thistransition radius to be approximately 4 km (total intrusion widthapproximately 8 km) for the laccoliths of the Henry Mountains.Thorey and Michaut (2014) provide a non-dimensionalized rela-tionship for the transition radius centered on the flexural parame-ter K (Turcotte et al., 1981). Elastic models suggest that craterswith radii C > 4K will display a flat floor and craters with C < 4Kwill display a domed floor. In Section 4 we will explore both thequalitative predictions from terrestrial studies and the quantitativeresults of Thorey and Michaut (2014) with regard to observed FFCmorphology.


4. Postulated effects of magmatic intrusion

The processes of magmatic intrusion and sill formationdescribed above, is predicted to be accompanied by specific typesof fracturing in the overlying strata. Additionally, the degassingof the magma could lead to venting or pyroclastic eruptions. Weexamine each of these processes as they could be expressed inthe setting of a floor-fractured crater, and compare observationswith our predictions in Section 5.

4.1. Fracture formation

The namesake characteristics of FFCs are floor fractures, whichcan be classified as concentric, radial, or polygonal (a combinationof concentric and radial) (Schultz, 1976; Jozwiak et al., 2012). Wepostulate that these fractures are formed as a direct result of craterfloor uplift as the sill/laccolith forms beneath the crater floor. Martiet al. (1994) performed experiments simulating magma reservoirinflation and deflation, and noted that in this setting, both concen-tric and radial fractures will form, with the concentric fracturesbeing focused around the regions of most extreme uplift. Thislocalization of fracturing and faulting above the regions of the sillthat experience the most extreme tilt was also discussed in theassessment of the Henry Mountain laccoliths by Jackson andPollard (1988). Jackson and Pollard (1988) also showed that forintrusions where the stress is concentrated at the periphery (tabu-lar sills), the concentric faulting was concentrated near the edgesand not the central parts of the intrusion. Studies of dome growthat Lascar Volcano in Chile have shown that, during periods of defla-tion, movement is accommodated along concentric fractures thatwere initially formed during the period of dome inflation(Matthews et al., 1996). Experiments on fractures associated withsalt-dome formation show that both concentric and radial frac-tures form early in the deformation process (Cloos, 1955).

Applying these observations to FFCs, we see that both radial andconcentric fractures are expected consequences of surface doming,such as the doming proposed by our model of magmatic intrusionand sill formation. Furthermore, we can postulate that in large cra-ters, with flat floors, uplift was likely to have been piston-like dueto a tabular intrusion, and this notion is supported tectonically forlarge craters by the observations of concentric fractures close tothe wall region (Figs. 3B, C, 7B, C, 8B and C), where the periphery

Fig. 14. Depiction of the major processes involved in the evolution and degassing of a sformation and settling, volatile exosolution and bubble rise, and possible associated sur

of the intrusion would have experienced the highest degree ofbending stresses. Conversely, in craters where the intrusion hasthe shape of a domed laccolith, the concentric fractures will be clo-ser to the center of the crater floor, and will mark the areas of mostextreme flexure. Smaller radial fractures form during extendedperiods of deformation and are thus expected in larger, flat-flooredcraters. We postulated that craters with large, central radial frac-tures experienced an initial period of centrally localized deforma-tion based on the experiments of Cloos (1955). Additionally, anysubsidence that occurs as a result of intrusion cooling will occuralong the existing concentric fractures; in the case of largeamounts of subsidence, these fractures may take on a fault/scarpmorphology (Matthews et al., 1996). Thorey and Michaut (2014)model the deformation of an elastic-sheet-like crater floor for thecase of floor-fractured craters. They similarly observe large, flat-floored craters with peripheral bending stresses and smallerdomed-floor craters with bending stresses focused toward the cra-ter interior. Their model parameters suggest, however, that theseeffects could also stem from the depth of the intrusion.

4.2. Magma degassing, venting, and pyroclastic eruptions

Another important implication of FFC formation by magmaticintrusion and sill formation is the subsequent cooling, evolution,and degassing of the sill or laccolithic magmatic intrusion. The var-ious consequences that can arise from this process in the context ofFFCs are illustrated in the schematic diagram in Fig. 14. The firstprocess to consider is the cooling and crystallization of the mag-matic intrusion. This problem is addressed after the manner ofTurcotte and Schubert (1982) as an application of the Stefan solu-tion and considers equal cooling rates from all edges; the intrusionis considered completely solidified at the point when the solidifica-tion boundary reaches the midpoint of the intrusion thickness. Wechose a representative intrusion thickness of 1 km. In this scenario,a 1 km thick intrusion takes approximately four years to solidify,although we emphasize this is only for solidification and not forcomplete cooling to the temperature of the surrounding countryrock. Also depicted in Fig. 14 is the process of crystal settlingwithin the intrusion as minerals crystallize. As suggested byMarsh (1989), there will exist a crystal capture front at the lowersolidification boundary, and a thinner capture front at the upper

ubcrater magmatic intrusion. Key processes include intrusion solidification, crystalface morphologies. Details of this evolution process are described in the text.

Table 2Material constants and properties.

Property Symbol Value

Fracture toughness K+a 100 MPa m1/2

Lunar gravity g 1.62 m s�2

Lunar crust density qc 2580 kg m�3

Lunar magma density qm 3220 kg m�3

Lunar breccia density qb 2410 kg m�3

Elastic modulus B 7.47 � 1010 PaMagma yield strength k 104 N m�2

Unit magma weight c 5215 kg m�2 s�2


solidification boundary, with crystals possibly being verticallytransported by bubble attachment.

The discussion of bubble attachment leads to one of the mostimportant consequences of magmatic intrusion evolution—magmadegassing/volatile production, and the possibility of associatedventing and pyroclastic eruptions. The primary volatile is CO whichis produced as a product of reduction reactions between Cgraphite

and Cr2O3, FeO, and Ti2O3 in the melt in concentrations of approx-imately 1000 ppm (Sato, 1979; Fogel and Rutherford, 1995). Wealso consider contributions from H2O, which could be present inthe melt in concentrations of 260–1410 ppm (Saal et al., 2008;Hauri et al., 2011). The production of CO is a pressure-dependentreaction, which is maximized at a pressure of 10 kPa, with smalleramounts of volatiles produced at lower pressures (Weitz et al.,1999). This pressure corresponds to a depth of�4 km on the Moon;thus, intrusions that are located at this depth are likely to producelarger amounts of CO, resulting in a higher gas pressure. Adaptingtreatments by Johnson and Pollard (1973) and Wichman andSchultz (1996), the depth of the intrusion, Te, can be related tothe intrusion radius, a, the country rock elastic modulus (hereYoung’s modulus with the assumption of m = 0.25), B, the magmaunit weight, cm, and the magma yield strength, k, by

T3e ¼

2ka þ cm

� �a4

5:33 � B : ð2Þ

Using the parameters listed in Table 2, this equation implies thatmaximum CO production occurs when a = �47 km, suggesting thatFFCs with D � 90–100 km are optimized for CO production withinthe magmatic intrusion.

We now turn to the possible results of this volatile production—intrusion degassing, either passively along fractures, or in activevulcanian eruptions. The former represents the case of magma(and volatiles) being transported directly along fractures fromthe intrusion to the surface, and is depicted in Fig. 14 as a surface

-3.5

-2.5

-1.5

-0.5

0 10 20 30 40

Rim Crest

Fracture

Domed Central Region

Ridge

Rim Crest

“V” Moat“V” Moat

A

A’

Diameter [km]

Dep

th [k

m]

A

Bohnenberger

Fig. 15. Craters (A) Bohnenberger (D = 33 km) (16.2�S, 40.0�E) and (B) Von Braun (D = 60deg. Note the overall domed floor profile of the crater Bohnenberger which is contraste

lava flow; the latter case is where the gas pressure within a par-tially filled fracture exceeds the failure criterion for the overlyingrocks, which is twice the tensile strength and �15 MPa(Touloukian et al., 1980; Tait et al., 1989; Head et al., 2002). Wewill now establish a set of conditions for which we would predict(1) surface lava flows, (2) venting, and (3) vulcanian eruptions.

(1) Surface lava flows: These are the result of direct magmatransport along fractures from the intrusion to the craterfloor. This depth of fracturing limits the region for expectedlava flows to the area adjacent to the crater wall, and theconcentric fractures associated with piston like uplift ofthe crater floor. Jackson and Pollard (1988) demonstrate thatdikes propagating from the intrusion are localized to thehighly bent, peripheral region of large, tabular intrusions.Thus, we would predict that lava flows may be present inlarge, flat-floored craters with wall-adjacent concentricfractures.

(2) Gas venting: This process could occur in a number of ways, assuggested by the geometry of Fig. 14 and the presence ofdark-halo craters in FFCs (Fig. 8). First, cooling and solidifica-tion at shallow depths will lead to pressure dependent vola-tile formation and magmatic foam buildup toward the top ofthe intrusion. A pure venting scenario does not necessarilyproduce a visibly dark deposit surrounding the vent inLROC-WAC images. The result of passive volatile leakageand degassing is broad floor subsidence, and perhaps locally,the drainage of overlying material to form a collapse crater.This latter type of feature forms primarily within existingfloor fractures that might have magma intruded beneaththem. The pressure in this gas + magma mixture is not highenough to fracture the overlying rock. Rather, we postulatethat the low pressure (�5 MPa) leads to unstable foamswhich drain and collapse (Jaupart and Vergniolle, 1989;Head et al., 2002), leaving a void space into which overlyingmaterial collapses and forms a vent (Wilson et al., 2011).This vent morphology would be predicted for host cratersof a moderate size, which have deep fractures (radial or con-centric), but are not large enough to allow for extensive COformation reactions and thus have a lower overall magmagas pressure.

(3) Vulcanian eruptions: Circular or elongate vent-like features,surrounded by a dark, mantling deposit (dark halo crater,DHC) have been documented in FFCs. In the crater Alphon-sus, Head and Wilson (1979) proposed that the DHCs wereformed by vulcanian eruptions. Specifically, they proposedthat magma rose to the level of the breccia lens (this same

-3

-2

-1

0

0 15 30 45 60 75

Rim CrestRim Crest

Wall Slumps

Fractures

A

A’

B Diameter [km]

Dep

th [k

m]

Von Braun

km) (41.1�N, 78.0�W), topographic profiles of the craters, LOLA gridded data 512 px/d to the comparatively flat floor profile of the crater Von Braun.

Fig. 16. Crater Alphonsus (D = 119 km) (13.4�S, 2.8�W), class 5 FFC, LROC-WACimage data. The white arrow indicates a vent located along one of the floor-fractures. There is no readily identifiable dark mantling material associated withthis vent, unlike the nearby vent and the larger easily identifiable dark halo craterson the floor of Alphonsus. Although the identified vent lacks visibly identifiabledeposits, spectral data reveals a small amount of pyroclastic material surroundingthis vent (Gaddis et al., 2011).


level representing surface extrusion in the adjacent MareNubium) beneath Alphonsus; then, localized volatile accu-mulations exceeded the tensile strength of the overlying cra-ter floor, resulting in localized explosive eruptions whichemplaced fine-grained juvenile and non-juvenile materialforming a DHC (Head and Wilson, 1979). We propose thatin the broad context of FFCs, these DHC morphologies formas magma partially fills fractures, or fractures form overareas of subsidiary diking from the main intrusion body.Subsequent gas build up in this magma produces a gas + foammixture with an overall pressure in excess of �15 MPa,resulting in fracture of the overlying material, and a vulca-

Fig. 17. Crater Alphonsus (D = 119 km) (13.4�S, 2.8�W), class 5 FFC. (A) Topographic colorAlphonsus crater sketch map, highlights the locations of dark halo crater pyroclastic depoconcentric fractures that are near to the crater wall. (For interpretation of the references t

nian eruption (Head and Wilson, 1979). Head et al. (2002)proposed this process and analyzed it in detail for the darkring deposit in Orientale Basin, and using the specific ventgeometry and deposit morphometry were able to placeconstraints on the mass and volume of the foam layer. Forthe purposes of our criterion we describe these qualitatively;however, we note that the analysis of Head et al. (2002) ispossible for any and all vulcanian eruption features in FFCs.We predict that these features will occur along fractures,specifically fractures near the periphery of the crater floor,where subsidiary diking from the intrusion is most likelyto occur; these features will also be more likely to occur inlarger craters that have a greater intrusion depth, and thusproduce more CO during reduction reactions, increasingthe available gas pressure needed to form these eruptions.

5. Discussion of predictions as applied to observed FFCmorphologies

In the previous section, we outlined predictions for the tectonicand volcanic effects of forming FFCs as a result of magmatic intru-sion and sill formation. We will now examine these tectonic andvolcanic predictions in the context of observed crater morpholo-gies, and then propose formation sequences for the craters Vitelloand Humboldt to test the predictions. Finally, we make predictionsabout the Bouguer gravity anomalies that might be expected forFFCs.

5.1. Fractures in FFCs

As described in Section 3.1, there exists a strong relationshipbetween the diameter of an uplifted crater floor region and thelocation of concentric fractures on the floor (Figs. 3–8). For largecraters with flat floors, bending stresses concentrate at the periph-ery of the intrusion, yielding concentric fractures close to the craterwall. This can be observed in the craters Humboldt (Fig. 3)D = 207 km, Gassendi (Fig. 5) D = 110 km, Von Braun (Fig. 7)D = 60 km, and Oppenheimer (Fig. 8) D = 208 km. As the diameterof the crater decreases, the intrusion becomes more centrallydomed, and the bending stresses move closer to the crater center(Jackson and Pollard, 1988), resulting in concentric fractures far-ther away from the crater wall, and coinciding with the regionsof greatest bending, as seen in craters Vitello (Fig. 4) D = 42 km

-coded LOLA altimetry (blue low, red high) overlain with LROC-WAC image data. (B)sits and the associated vents. The vents are located along fractures, and in particularo color in this figure legend, the reader is referred to the web version of this article.)

A

C

B

Fig. 18. Evolution of Vitello crater by magmatic intrusion and sill formation(D = 44 km) (30.4�S, 37.5�W). (A) Initial unmodified crater profile. (B) Verticallypropagating dike is halted by underdense brecciated region beneath crater. Dikepropagates laterally to edges of crater floor, forming sill. (C) Sill inflates, producing alaccolith which domes and deforms the overlying crater floor. Bending stresses arehighest in deformed region, and manifest as concentric fractures around theuplifted central crater floor.

A

D

C

B

Fig. 19. Evolution of Oppenheimer crater by magmatic intrusion and sill formation(D = 208 km) (35.2�S, 166.3�W). (A) Initially unmodified crater profile, Oppenhei-mer and associated craters Oppenheimer H and Oppenheimer U. (B) Verticallypropagating dike stalls beneath crater, then propagates laterally beneath the crater,halting at crater wall region. (C) Large intrusion diameter relative to intrusion depthconcentrates bending stresses at the periphery of the intrusion, where faulting anduplift of the crater floor then occurs (faults marked by arrows). Overall shallowingand flattening of the crater floor. (D) Magma from the intrusion is transported alongthe concentric faults and fractures. In Oppenheimer H, magma reaches the surfaceand lava flows form on the crater floor. In localized regions, gas overpressurizationinduces fracturing, resulting in deposition of pyroclastic material.


and Bohnenberger (Fig. 6) D = 33 km. Radial fractures emanatefrom regions of strong uplift, similar to concentric fractures; how-ever, radial fractures are created in response to uplift in the centerof the crater floor, as opposed to the bending stresses from theedges of the intrusion. Examples of these fractures can be seen incraters Bohnenberger (Figs. 6, 13A, and 15), Von Braun (Figs. 7,13B, and 15), and Humboldt (Fig. 3).

5.2. Evidence of magma degassing in FFCs

In Section 3.2 we outlined three possible surface manifestationsof degassing of the magmatic intrusion: (1) surface lava flows, (2)vent formation via foam collapse, and (3) vulcanian eruptions withassociated pyroclastic deposits. Surface lava flows are clearly seenin the crater Humboldt (Fig. 3), where they are manifested as cres-cent-shaped patches of material adjacent to the crater wall. Addi-tionally, it is notable that two of the deposits in the westernportion of the crater appear to mantle a large concentric fracturein this region. This suggests that the fracture served as a pathway

for magma transport from the sub-crater floor intrusion to the sur-face of the crater floor, which is consistent with our previousanalysis.

Table 3Morphometric calculation of intrusion thickness and volume.

Crater Average crater depth (km) Comparable fresh crater depth (km) Intrusion thickness (km) Approximate intrusion volume (km3)

Gassendi 2.4 4.3 1.9 2 � 104

Vitello 1.6 3.2 1.6 1 � 104

Humboldt 2.9 5.2 2.3 8 � 104


In the investigation of both vent formation and vulcanian erup-tions, we emphasize that the process has the ability to occur fromall intrusions; however, the individual geometry of the intrusionand the fractures leading to the crater floor will determine whetheror not the process occurs. The second case, of vent formation inmoderately-sized craters, pertains to only a small subset of craterswith vents, specifically those craters with vents lacking readilyidentifiable dark mantling material; an example of this can be seenin a recently identified vent in crater Alphonsus (Fig. 16),D = 119 km (Gaddis et al., 2011). Although Alphonsus hosts numer-ous vents with visibly identifiable dark halos, the pyroclastic signa-ture around this vent is only discernable using spectral data. Thelack of a large volume of material ejected from the vent suggeststhat there was passive degassing of a magmatic foam, leading toa volume reduction and collapse to form the vent (Fig. 14). Differ-ing from these collapse-crater vents, sites of postulated vulcanianeruptions are marked by associated dark mantling deposits and,most importantly, spectrally distinct pyroclastic deposits. The keydifference in the formation of these features is the increased over-pressure in the magmatic foam, resulting in a higher volume, activedegassing beneath the vent, erupting a larger volume of materialresulting in a dark halo deposit. These deposits should be associ-ated with vents and fractures on the crater floor, and should bepredominately in larger craters (D > 50 km). Gaddis et al. (2003)used Clementine data to map all spectrally identifiable pyroclasticdeposits on the Moon, and their catalog contained manyrecognized FFCs: Daniell, Gaudibert, Airy, Taruntius, Franklin,Doppelmeyer, Lavoisier, Mersenius, Atlas, Schlüter, Alphonsus,Cleomedes, Petavius, Gauss, Humboldt, Oppenheimer, andSchrödinger (Jozwiak et al., 2012). Excepting the craters Daniell,Gaudibert and Airy, all of the craters with identified pyroclasticdeposits are larger than 50 km in diameter. The craters Daniell,Gaudibert, and Airy do not have readily identifiable vents, thoughall three craters have significantly uplifted and deformed craterfloors. Additionally, all deposits are located near fractures, and inmost cases have associated vents, as seen in Alphonsus (Fig. 17)and Oppenheimer (Fig. 8). Craters of this diameter have postulatedintrusion depths coincident with the optimal volatile productiondepth, further strengthening the hypothesis that these observedpyroclastic deposits are the result of vulcanian eruptions fromthe magmatic intrusion.

5.3. Formation sequences for different types of floor-fractured craters:Vitello and Oppenheimer

The craters Vitello and Oppenheimer represent two distinctlydifferent FFC morphologies: Vitello (Fig. 4) is a domed-floor Class2 FFC with strong concentric fractures (Jozwiak et al., 2012), andOppenheimer (Fig. 8) is a large, flat-floored Class 6 FFC withwall-adjacent concentric fractures, vents and pyroclastic deposits,and is host to additional smaller FFCs (Jozwiak et al., 2012).

The modification sequence for the crater Vitello begins with anunmodified fresh crater profile, shown in Fig. 18A. A verticallypropagating dike beneath the crater encounters the breccia lensand, upon stalling in the highly fractured subcrater breccia region,the magma propagates laterally to the edges of the crater floor,forming a sill (Fig. 18B). During the ensuing stage, magma

continues to inflate the sill, forming a laccolith; the bending ofthe crater floor in response to the uplift creates concentric frac-tures, localized over the zones of greatest bending (Fig. 18C; com-pare with Fig. 4). This process, results in the two most importantmorphologic features of the crater Vitello—the domed floor, andthe concentric fractures surrounding the uplifted region (Fig. 4).

The modification of the crater Oppenheimer begins with anunmodified crater profile for Oppenheimer itself and the two smal-ler craters Oppenheimer U and Oppenheimer H, shown in cross-section in Fig. 19A. The process begins when an ascending dikestalls in the highly fractured subcrater breccias, and then propa-gates laterally, forming a sill. As magma inflates the sill, bendingstresses concentrate at the periphery of the sill (Fig. 19B), as aresult of the large ratio of crater floor diameter to intrusion depth.The peripheral bending stresses exceed the yield strength of theoverlying rocks, and fracturing and uplift of the crater floor occurs,manifested on the crater floor as concentric fractures close to thecrater wall, and an overall flat crater floor (Fig. 19C; compare withFig. 8). Magma from the intrusion may rise into these fractures.Beneath craters Oppenheimer U and Oppenheimer H, this frac-ture-transported magma encounters additionally brecciated rocks,stalls, forms a sill, and uplifts the floors of these craters, yieldingsmall FFCs within a larger FFC (Fig. 19D; compare with Fig. 8). Thisprocess of subsidiary diking and magma transport in fractures atthe top of the laccolith is also observed in terrestrial laccoliths(Corry, 1988). There also exist regions where magma ascended intofractures, but did not propagate to the surface before solidifying;here exsolved volatiles from the magma built up until theyexceeded the tensile strength of the overlying rock, and thenerupted in a vulcanian explosion marked by pyroclastic depositsaround vents (e.g. Head and Wilson, 1979; Head et al., 2002)(Fig. 19D).

6. Implications for gravity anomalies and future work

High-resolution gravity data would permit one to develop fur-ther tests to determine whether specific floor-fractured cratersare the result of magmatic intrusion and sill formation, or viscousrelaxation. Geophysical studies of terrestrial impact craters showthat craters are associated with negative Bouguer anomalies(Dabizha and Fedynsky, 1975; Pilkington and Grieve, 1992).Dvorak and Phillips (1977) showed that young Copernican/Eratos-thenian-aged lunar craters also exhibited negative Bouguer anom-alies, with the implication that the negative anomaly is the resultof the low density, brecciated region beneath the crater. However,Dvorak and Phillips (1978) analyzed several Imbrian-aged craters,including the recognized FFCs Petavius and Humboldt, andreported that these older craters had a net zero Bouguer anomaly.They concluded that either compaction of the low-density brecciamaterial or the addition of a higher density component is neces-sary to yield a net zero Bouguer anomaly in these older craters. Aformation mechanism of magmatic intrusion and sill formationfor FFCs would be one such way to introduce a high density com-ponent beneath the crater (Schultz, 1976; Dvorak and Phillips,1978; Wichman and Schultz, 1995). Dvorak and Phillips (1978)analyzed the process of sill formation as a means of creating theobserved Bouguer anomaly, and determined that it was an unlikely


source because of predicted intrusion dimensions and the uncer-tain mechanics of halting the intruding magma beneath the crater.

In Section 3 of this paper, we showed that the density of thelunar crust aided by the further decreased density of the breccialens material is sufficient to halt vertical propagation of an over-pressurized, negatively buoyant dike, and allow for sill formationbeneath the crater. Thus, sill/laccolith intrusion is a very plausiblemechanism for providing shallow, high density slabs below craterfloors. Furthermore, the morphometric data from LOLA altimetryprovides an estimate of the deviation of the crater floor depth fromthat of a fresh crater (Jozwiak et al., 2012). This information,together with the slope of the crater floor, provides an independentestimate of the radius and thickness of the slab. Thus this informa-tion (Table 3) can be used in concert with data from the GRAIL(Gravity Recovery and Interior Laboratory) mission (Zuber et al.,2013) (e.g. Jozwiak et al., 2014). For example, using a simple Bou-guer slab geometry, for the crater Gassendi, with an intrusionthickness of 2 km (Table 3), a lunar magma density of 3000 kg/m3, and a lunar crustal density of 2550 kg/m3, we would predicta Bouguer gravity anomaly of �40 mGal. This magnitude signalwill be readily detectable in the GRAIL data.

Jozwiak et al. (2014) found that a majority (80%) of the analyzedFFCs had a Bouguer gravity anomaly that was positive relative tothe surrounding region, and the remaining craters that lacked adistinct signal were predominately smaller craters with floordimensions close to the spatial resolution of the gravity model.Thorey et al. (2014) carried out a statistical survey of the lunarFFC population presented by Jozwiak et al. (2012), comparing theBouguer gravity anomalies of FFCs to the Bouguer anomalies ofrandomly sampled lunar complex craters. The work found weaksupport for positive Bouguer gravity anomalies within FFCs. How-ever, this work included several subclasses of FFCs which wereexcluded from the Jozwiak et al. (2014) study on the basis of cratersize or location, possibly introducing more scatter and weaker rela-tionships into the results.

We are now pursuing a detailed analysis of these gravity data,and the implications for both FFC formation and crater structure(Jozwiak et al., 2014). We will use detailed crater morphometryto determine intrusion dimensions and the predicted contributionto the Bouguer gravity anomaly. Table 3 shows these morphomet-ric calculations for three craters. These results can then be com-pared with the observed Bouguer anomaly to gain insight intothe subcrater intrusion environment, and the overall subcraterstructure.

7. Conclusions

We outline morphometric and morphological support for thehypothesis that lunar FFCs formed as a result of magmatic intru-sion and sill formation (Jozwiak et al., 2012). We find that the dis-tribution and characteristics of the FFC population correlatesstrongly with crustal thickness and the predicted frequencydistribution of overpressurization values of magmatic dikes. For atypical nearside lunar crustal thickness, dikes with high overpres-surization values favor surface eruptions, medium values favorintrusion and sill formation, and low values favor formation ofsolidified dikes concentrated lower in the crust. We have outlinedand analyzed the mechanics of this process: (1) a dike stalls in theunderdense breccia region beneath a lunar crater; (2) the drivingpressure exceeds the local lithostatic pressure promoting lateralpropagation and sill formation; (3) the larger overburden pressurefrom the crater rim and crater wall halt propagation at the edges ofthe crater floor; and (4) the sill grows to form a laccolith, with theratio of the crater diameter relative to the intrusion depth govern-ing the overall morphology of the laccolith. We then inferred

morphologic and geophysical consequences associated with themagmatic intrusion formation hypothesis, and compared thesepredictions with observed crater morphologies. The process ofuplifting the crater floor produces fractures on the floor, witharcuate, concentric fractures being the most prominent. These con-centric fractures should localize over areas of highest bendingstress—near the crater wall for large craters with piston-upliftedfloors, and interior to the crater over the domed floor sections forsmaller craters. These predicted patterns are observed and are con-sistent with FFC morphologies. We proposed three potential out-comes of intrusion evolution including surface lava flows adjacentto the crater wall, vents along fractures with no associated darkmantle deposits, and vents along fractures with associated darkmantle pyroclastic deposits. All three of these predicted morphol-ogies are observed in FFCs, and the specific expressions of thesephenomena are governed by the individual subcrater/ intrusionenvironments. Thus, on the basis of the strong morphologic sup-port for the predicted outcomes of our model, we conclude thatthe majority of lunar FFCs are the result of crater modification inresponse to magmatic intrusion and sill formation beneath thecrater. The process of magmatic intrusion and sill formation alsopresents a strong geophysical constraint in that the dense magmabody is predicted to provide a positive Bouguer anomaly contribu-tion, altering the normally negative Bouguer anomaly of impactcraters. We make predictions that can be tested using GRAIL grav-ity data, and that will provide further tests for the magmatic originof floor-fractured craters.

Acknowledgments

We gratefully acknowledge the support of NASA Harriet G. Jen-kins Fellowship (Grant NNX13AR86H) to L.M. Jozwiak. We alsogratefully acknowledge financial support from the NASA LunarReconnaissance Orbiter (LRO) Mission, Lunar Orbiter LaserAltimeter (LOLA) Experiment Team (Grants NNX11AK29G andNNX13AO77G), the NASA Gravity Recovery and Interior Laboratory(GRAIL) Mission Guest Scientist Program (Grant NNX12AL07G) andthe NASA Solar System Exploration Research Virtual Institute(SSERVI) grant for Evolution and Environment of ExplorationDestinations under cooperative agreement number NNA14AB01Aat Brown University. Additionally, we acknowledge Jay Dicksonfor his invaluable assistance in data processing.

References

Baker, D.M.H. et al., 2011. The transition from complex crater to peak-ring basin onthe Moon: New observations from the Lunar Orbiter Laser Altimeter (LOLA)instrument. Icarus 214, 377–393.

Baker, D.M.H., Head III, J.W., Neumann, G.A., Smith, D.E., Zuber, M.T., 2012. Thetransition from complex crater to multi-ringed basins on the Moon:Quantitative geometric properties from Lunar Reconnaissance Orbiter LunarOrbiter Laser Altimeter (LOLA) data. J. Geophys. Res. 117, E00H16.

Brennan, W.J., 1975. Modification of premare impact craters by volcanism andtectonism. The Moon 12, 449–461.

Cathles, L.M., 1975. The Viscosity of the Earth’s Mantle. Princeton University Press,Princeton, NJ, pp. 386.

Cloos, E., 1955. Experimental analysis of fracture patterns. Bull. Geol. Soc. Am. 66,241–266.

Corry, C.E., 1988. Laccoliths; mechanics of emplacement and growth. GSA SpecialPaper 220.

Dabizha, A.I., Fedynsky, V.V., 1975. The Earth’s ‘‘star wounds’’ and their diagnosis bygeophysical methods. Zamlya Vselennaya 3, 56–64 (in Russian).

Daneš, Z.F., 1965. Rebound processes in large craters. Astrogeologic Studies, Annu.Prog. Rep. A. U.S. Geol. Surv., Washington, DC, pp. 81–100.

Dence, M.R., Grieve, R.A.F., Robertson, P.B., 1977. Terrestrial impact structures –Principle characteristics and energy considerations. In: Roddy, D.J., Pepin, R.O.,Merrill, R.B. (Eds.), Impact and Explosion Cratering. Pergamon Inc., New York,pp. 247–275.

Dombard, A.J., Gillis, J.J., 2001. Testing the viability of topographic relaxation as amechanism of the formation of lunar floor-fractured craters. J. Geophys. Res.106 (E11), 27901–27909.

http://refhub.elsevier.com/S0019-1035(14)00633-2/h0005























Dvorak, J.J., Phillips, R.J., 1977. The nature of the gravity anomalies associated withlarge young lunar craters. Geophys. Res. Lett. 4, 380–382.

Dvorak, J.J., Phillips, R.J., 1978. Lunar Bouguer gravity anomalies: Imbrian agecraters. Proc. Lunar Sci. Conf. 7, 3651–3668.

Elkins-Tanton, L.T., Hager, B.H., Grove, T.L., 2004. Magmatic effects of the lunar lateheavy bombardment. Earth Planet. Sci. Lett. 222, 17–27.

Fogel, R.A., Rutherford, M.J., 1995. Magmatic volatiles in primitive lunar glasses: I.FTIR and EPMA analyses of Apollo 15 green and yellow glasses and revision ofthe volatile-assisted fire-fountain theory. Geochim. Cosmochim. Acta 59, 201–215.

Gaddis, L.R., Staid, M.I., Tyburczy, J.A., Hawke, B.R., Petro, N.E., 2003. Compositionalanalyses of lunar pyroclastic deposits. Icarus 161, 262–280.

Gaddis, L.R., Klem, S., Gustafson III, J.O., Hawke, B.R., Giguere, T.A., 2011. Alphonsusdark-halo craters: Identification of additional volcanic vents. Lunar Planet. Sci.XLII. Abstract #2691.

Gaddis, L.R. et al., 2013. ‘‘New’’ volcanic features in lunar, floor-fracturedOppenheimer crater. Lunar Planet. Sci. XLIV. Abstract #2262.

Hall, J.L., Solomon, S.C., Head, J.W., 1981. Lunar floor-fractured craters: Evidence ofviscous relaxation of crater topography. J. Geophys. Res. 86, 9537–9552.

Hauri, E.H., Weinreich, T., Saal, A.E., Rutherford, M.C., Van Orman, J.A., 2011. Highpre-eruptive water contents preserved in lunar melt inclusions. Science 333(6039), 213–215.

Hawke, B.R. et al., 2004. The origin of lunar crater rays. Icarus 170, 1–16.Hawke, B.R. et al., 2006. The composition and origin of lunar crater rays:

Implications for the Copernican–Eratosthenian boundary. Lunar Planet. Sci.XXXVII. Abstract #1133.

Head, J.W., Wilson, L., 1979. Alphonsus-type dark-halo craters: Morphology,morphometry, and eruption conditions. Proc. Lunar Sci. Conf. X, 525–527.

Head III, J.W., Wilson, L., 1989. Basaltic pyroclastic eruptions: Influence of gas-release patterns and volume fluxes on fountain structure, and the formation ofcinder cones, splatter cones, rootless flows, lava ponds and lava flows. J.Volcanol. Geotherm. Res. 37, 261–271.

Head, J.W., Wilson, L., 1992. Lunar mare volcanism: Stratigraphy, eruptionconditions, and the evolution of secondary crusts. Geochim. Cosmochim. Acta56, 2155–2175.

Head, J.W., Wilson, L., Weitz, C.M., 2002. Dark ring in southwestern Orientale Basin:Origin as a single pyroclastic eruption. J. Geophys. Res. 107 (E1), 5001.

Head, J.W. et al., 2010. Global distribution of large lunar craters: Implications forresurfacing and impactor populations. Science 329, 1504–1507.

Hiesinger, H., Head, J.W., 2006. New views of lunar geoscience: And introductionand overview. Rev. Mineral. Geochem. 60 (1), 1–81.

Hiesinger, H., Head III, J.W., Wolf, U., Jaumann, R., Neukum, G., 2011. Ages andstratigraphy of lunar mare basalts: A synthesis. Geol. Soc. America Special Paper477, 1–51.

Hiesinger, H. et al., 2012. How old are young lunar craters? J. Geophys. Res. 117,E12.

Hurwitz, D.M., Head, J.W., Hiesinger, H., 2013. Lunar sinuous rilles: Distribution,characteristics, and implications for their origin. Planet. Space Sci. 79–80, 1–38.

Jackson, M.D., Pollard, D.D., 1988. The laccolith-stock controversy: New results fromthe southern Henry Mountains, Utah. Geol. Soc. Am. Bull. 100, 117–139.

Jaupart, C., Vergniolle, S., 1989. The generation and collapse of a foam layer at theroof of a basaltic magma chamber. J. Fluid Mech. 203, 347–380.

Johnson, A.M., Pollard, D.D., 1973. Mechanics of growth of some laccolithicintrusions in the Henry Mountains, Utah. I. Field observations, Gilbert’smodel, physical properties and flow of magma. Tectonophysics 18, 261–309.

Jozwiak, L.M., Head, J.W., Zuber, M.T., Smith, D.E., Neuman, G.A., 2012. Lunar floor-fractured craters: Classification, distribution, origin and implications formagmatism and shallow crustal structure. J. Geophys. Res. 117, E11.

Jozwiak, L.M. et al., 2014. Lunar floor-fractured craters: Intrusion emplacement andassociated gravity anomalies. Lunar Planet. Sci. XLV. Abstract #1464.

Kiefer, W.S., Macke, R.J., Britt, D.T., Irving, A.J., Consolmagno, G.J., 2012. The densityand porosity of lunar rocks. Geophys. Res. Lett. 39, L07201.

Kramer, G.Y., Jolliff, B.L., Neal, C.R., 2008. Distinguishing high alumina mare basaltsusing Clementine UVVIS and Lunar Prospector GRS data: Mare Moscoviense andMare Nectaris. J. Geophys. Res. 113, E01002.

Losiak, A. et al., 2009. A new lunar impact crater database. Lunar Planet. Sci. XL.Abstract #1532.

Marsh, B.D., 1989. Magma chambers. Annu. Rev. Earth Planet. Sci. 17, 439–474.Marti, J., Albay, G.J., Redshaw, L.T., Sparks, R.S.J., 1994. Experimental studies of

collapse calderas. J. Geol. Soc. Lond. 151, 919–929.Masursky, H., 1964. A preliminary report on the role of isostatic rebound in the

geologic development of the lunar crater Ptolemaeus. Astrogeologic Studies,Annu. Prog. Rep. A, U.S. Geol. Surv., Washington, DC, pp. 102–134.

Matthews, S.J., Gardeweg, M.C., Sparks, R.S.J., 1996. The 1984 to 1996 cyclic activityof Lascar Volcano, northern Chile: Cycles of dome growth, dome subsidence,degassing and explosive eruptions. Bull. Volcanol. 59, 72–82.

Michaut, C., 2011. Dynamics of magmatic intrusions in the upper crust: Theory andapplications to laccoliths on Earth and the Moon. J. Geophys. Res. 116, B05205.

Moore, H.J., 1976. Missile impact craters (White Sands Missile Range, New Mexico)and applications to lunar research. U.S. Geol. Surv. Prof. Paper 812-B. p. 47.

Parfitt, E.E., 1991. The role of rift-zone storage in controlling the site and timing oferuptions and intrusions of Kilauea Volcano, Hawaii. J. Geophys. Res. 96 (B6),10101–10112.

Pike, R.J., 1980. Geometric interpretation of lunar craters. U.S. Geol. Surv. Prof.Paper, 1046-C, pp. C1–C77.

Pilkington, M., Grieve, R.A.F., 1992. The geophysical signature of terrestrial impactcraters. Rev. Geophys. 30 (2), 161–181.

Pohl, J., Stöffler, D., Gall, H., Ernston, K., 1977. The Reis impact crater. In: Roddy, D.J.,Pepin, R.O., Merrill, R.B. (Eds.), Impact and Explosion Cratering. Pergamon Inc.,New York, pp. 343–404.

Rubin, A.M., Pollard, D.D., 1987. Origins of blade-like dikes in volcanic rift zones. In:Volcanism in Hawaii. U.S. Geol. Soc. Prof. Paper 1350, pp. 1449–1470.

Saal, A.E., Hauri, E.H., Cascio, M.L., Van Orman, J.A., Rutherford, M.C., Cooper, R.F.,2008. Volatile content of lunar volcanic glasses and the presence of water in theMoon’s interior. Nature 454, 192–196.

Sato, M., 1979. The driving mechanism of lunar pyroclastic eruptions inferred fromthe oxygen fugacity behavior of Apollo 17 orange glass. Proc. Lunar Sci. Conf. X,311–325.

Schultz, P.H., 1976. Floor-fractured lunar craters. The Moon 15, 241–273.Solomon, S.C., Head, J.W., 1980. Lunar mascon basins: Lava filling, tectonics, and

evolution of the lithosphere. Rev. Geophys. Space Phys. 18, 107–141.Tait, S., Jaupert, C., Vergniolle, S., 1989. Pressure, gas content, and eruptive

periodicity of a shallow crystalizing magma chamber. Earth Planet. Sci. Lett.92, 107–123.

Thaisen, K.G. et al., 2011. Geology of the Moscoviense basin. J. Geophys. Res. 116,E00G07.

Thorey, C., Michaut, C., 2014. A model for the dynamics of crater-centered intrusion:Application to lunar floor-fractured craters. J. Geophys. Res.: Planets. http://dx.doi.org/10.1002/2013JE004467.

Thorey, C., Michaut, C., Wieczorek., M. 2014. Gravitational signatures of lunar floorfractured craters. LPSC XLV. Abstract #2225.

Touloukian, Y.S., Judd, W.R., Roy, R.F., 1980. Physical Properties of Rocks andMinerals. McGraw-Hill, New York.

Turcotte, D.L., Schubert, G., 1982. Geodynamics Applications of Continuum Physicsto Geological Problems. Wiley, New York.

Turcotte, D.L., Willemann, R.J., Haxby, W.F., Norberry, J., 1981. Role of membranestresses in the support of planetary topography. J. Geophys. Res. 86 (B5), 3951–3959.

Vaughan, W.M., Head III, J.W., 2014. Impact melt differentiation in the South Pole-Aitken basin: Some observations and speculations. Planet. Space Sci. 91, 101–106.

Vaughan, W.M., Head III, J.W., Wilson, L., Hess, P.C., 2013. Geology and petrology ofenormous volumes of impact melt on the Moon: A case study of the OrientaleBasin impact melt sea. Icarus 223, 749–765.

Weitz, C.M., Rutherford, M.J., Head, J.W., 1999. Ascent and eruption of a lunar high-titanium magma as inferred from the petrology of the 74001/2 drill core.Meteorit. Planet. Sci. 34, 527–540.

Wichman, R.W., Schultz, P.H., 1995. Floor-fractured craters in Mare Smythii andwest of Oceanus Procellarum: Implications of crater modification by viscousrelaxation and igneous intrusion models. J. Geophys. Res. 100 (E10), 21201–21218.

Wichman, R.W., Schultz, P.H., 1996. Crater-centered laccoliths on the Moon:Modeling intrusion depth and magmatic pressure at the crater Taruntius. Icarus122, 193–199.

Wieczorek, M.A. et al., 2013. The crust of the Moon as seen by GRAIL. Science 339,671–675.

Wilhelms, D.E., 1987. The geological history of the Moon. USGS Prof. Paper 1348.Wilson, L., Head, J.W., 1981. Ascent and eruption of basaltic magma on the Earth

and Moon. J. Geophys. Res. 84 (B4), 2971–3001.Wilson, L., Hawke, B.R., Giguere, T.A., Petrycki, E.R., 2011. An igneous origin for Rima

Hyginus and Hyginus crater on the Moon. Icarus 215, 584–595.Wilson, L., Head, J.W., Tye, A.R., 2014. Lunar regional pyroclastic deposits: Evidence

for eruption from dikes emplaced into the near-surface crust. Lunar Planet. Sci.XLV. Abstract #1223.

Yingst, R.A., Head, J.W., 1997. Volumes of lunar lava ponds in South Pole-Aitken andOrientale Basins: Implications for eruption conditions, transport mechanismsand magma source regions. J. Geophys. Res. 102, 10909–10931.

Yingst, R.A., Head III, J.W., 1999. Geology of mare deposits in South Pole-Aitkenbasin as seen by Clementine UV/VIS data. J. Geophys. Res. 104 (E8), 18957–18979.

Zuber, M.T. et al., 2013. Gravity field of the Moon from the Gravity Recovery andInterior Laboratory (GRAIL) Mission. Science 339 (6120), 668–671.






















































































http://dx.doi.org/10.1002/2013JE004467

http://dx.doi.org/10.1002/2013JE004467






































Documents

Brown University Planetary Geosciences - Lunar …Lunar ﬂoor-fractured craters as magmatic intrusions: Geometry, modes of emplacement, associated tectonic and volcanic features,