Marine facies of carbonate and siliciclastic sediments deposited on top of the upper Devonian Alamo Breccia Member identify the shape and size of the Alamo impact crater in south-central Nevada (western USA). There are 13 measured sections that record peritidal to deep-subtidal deposition across the impacted platform, and these are correlated to three regional depositional sequences above the Alamo Breccia Member. Facies and accommodation patterns identify a concave seafloor that we interpret as the post-impact legacy of the Alamo crater. Together with isopach and lithostratigraphic trends in the underlying Alamo Breccia Member, a new map of the Alamo crater is presented showing the eastern outer rim fault and the annular trough. Size estimates were made using the newly defined crater features and linear scaling relationships from other marine-target complex craters. Revised dimensions of the Alamo crater place its transient diameter between 37 and 65 km, and its apparent diameter between 111 and 150 km. These estimates are more than double previous estimates based on the biostratigraphy of the Alamo Breccia Member. If correct, these new estimates place the Alamo crater as one of the largest marine impacts of the Phanerozoic, and conservatively larger than the well-studied Eocene Chesapeake Bay crater.


Marine bolide impact events are underrepresented in the rock record due to their low preservation potential. Consequently, few studies exist documenting marine impact crater size, morphology, and effects on sedimentation patterns; of the 27 known marine impact craters on Earth, 20 of them are currently located on land (Dypvik and Jansa, 2003). The Late Devonian Alamo impact of south-central Nevada (western USA) is one such case, providing a rare opportunity to study a marine impact in outcrop at the regional scale.

The Alamo impact occurred on a carbonate platform along the western margin of North America. This catastrophic event is now expressed in the Guilmette Formation (Late Devonian, Frasnian) across present-day south-central Nevada and western Utah (Fig. 1). The resultant impact stratum, known as the Alamo Breccia Member, covers an area of ∼28,000 km2 and is one of the largest and best-exposed marine impact deposits on Earth (Pinto and Warme, 2008). Evidence supporting an impact origin for the Alamo event includes melt breccia, shatter cone–like structures, carbonate accretionary lapilli, iridium anomalies, and shocked quartz (Pinto and Warme, 2008).

Post-impact tectonism throughout the region has obscured the original crater morphology and buried important strata, making it difficult to correlate between sections and characterize the impact crater (Pinto and Warme, 2008). Prior descriptions of the impact crater are based on the lithostratigraphy and features of the Alamo Breccia Member (Warme and Sandberg, 1995; Pinto and Warme, 2008). While useful for regional correlation, this terminology does not relate the deposits to a complex crater model as expected for marine bolide impacts (Dypvik and Jansa, 2003; Dypvik and Kalleson, 2010). Estimates of the final crater diameter have relied exclusively on the extent and composition of the Alamo Breccia Member and not on geomorphic features specific to marine impact craters.

The aim of this paper is to interpret crater morphology based on post-impact depositional environments in the context of a regional sequence stratigraphic framework. By identifying key boundaries of the crater margin, we make new size estimates of the Alamo crater based on linear scaling relationships from well-studied seismically imaged marine impact craters (Melosh, 1989; Dypvik and Kalleson, 2010). These methods could prove useful in estimating the size of other marine impact craters that lack seismic data, or that are associated with post-impact tectonism that has obscured the original crater morphology.



The Alamo impact occurred ca. 382 Ma on a shallow-marine, west-facing carbonate platform during deposition of the Guilmette Formation (Sandberg and Morrow, 1998). Three members compose the Guilmette Formation: the lower member, the Alamo Breccia Member, and the upper member (Fig. 2) (Ackman, 1991).

The lower member consists of a basal yellow slope-forming interval capped by a ledge-forming interval (Fig. 2) (Ackman, 1991; Sandberg et al., 1997). The yellow slope-forming interval comprises thinly bedded silty dolostone, and its base is marked by beds of digitate stromatolites (Sandberg et al., 1997). The ledge-forming interval consists of ∼100 m of intertidal and subtidal mudstones that transition eastward to thin, silty dolostone and microbial laminite (Sandberg et al., 1997). The Alamo bolide struck the ledge-forming interval of the lower member.

The Alamo Breccia Member is a carbonate megabreccia that forms the regional middle member of the Guilmette Formation (Fig. 2) (Pinto and Warme, 2008). It unconformably overlies the ledge-forming interval of the lower member (Warme et al., 1991). The base of the Alamo Breccia Member (D unit; Warme and Sandberg, 1995) is marked by a monomict, fluidized detachment breccia typically <3 m thick. Above this are tilted megaclasts (to 80 m thick) of pre-impact carbonate platform rocks (C unit; Warme and Sandberg, 1995) from the lower member of the Guilmette Formation; these are overlain by 1–30-m-thick surge and resurge breccia deposits (B and A units; Warme and Sandberg, 1995). These deposits grade normally upsection into sand- and mud-sized particles, marking the top of strata deposited during the Alamo impact. The top of the Alamo Breccia Member is here interpreted as an isochronous datum. Excavation of the lower member of the Guilmette Formation and emplacement of the Alamo Breccia Member within the impact crater may have resulted in a net loss or gain of sediment thickness, affecting regional accommodation (Fig. 3).

The upper member was deposited after the Alamo impact. It is composed of subtidal limestone, dolomite, and quartz sandstone (Fig. 2) (Sandberg et al., 1997; Chamberlain, 1999). Stromatoporoid reefs and mud mounds are locally present within the upper member throughout the impact region (Dunn, 1979; Tapanila and Ekdale, 2004; Tapanila et al., 2014). Quartzarenite is present ∼40–100 m upsection from the top of the Alamo Breccia Member and represents a shift from carbonate- to siliciclastic-dominated deposition.

Post-Impact Tectonic History

The deformation history of the Alamo crater during Paleozoic time is unknown. No faulting with appreciable displacement is associated with the Antler orogeny (Late Devonian to Mississippian). However, during Permian through Cretaceous time, the central Nevada thrust belt developed in Lincoln County within the hinterland of the Sevier orogenic belt (Misch, 1960). This produced parallel sets of north-striking thrust faults as part of the Garden Valley thrust system (Tschanz and Pampeyan, 1970; Bartley and Gleason, 1990; Armstrong, 1991; Taylor et al., 2000). To the east, the Mount Irish thrust includes three correlated faults (Golden Gate–Mount Irish–East Pahranagat), and to the west, the Rimrock, Freiberg, and Lincoln thrusts are correlated southward out of the field area. All localities in this study, except Tempiute Mountain, are present within the hanging wall or footwall of the Mount Irish thrust (Fig. 4). East-directed compression on the Mount Irish thrust is likely <10 km, with an overall increase in magnitude southward to the East Pahranagat fault (Taylor et al., 2000).

The Tempiute Mountain section is located on the west side of the Monte Mountain, Lincoln, and Schofield Pass thrusts, suggesting that it may have moved significantly eastward by as much as 24–30 km (Taylor et al., 1994). The Pahranagat shear zone accounts for 9–16 km of sinistral slip (Tschanz and Pampeyan, 1970; Liggett and Ehrenspeck, 1974; Ekren et al., 1977). Other faulting events in the region are of a much smaller scale, including normal and oblique faulting (<1 km slip) associated with the passage of Tertiary volcanic centers (e.g., Taylor and Switzer, 2001); the east-dipping Seaman breakaway fault (2 km slip; Taylor and Bartley, 1992); and high-angle normal faulting of Basin and Range extension. Greater magnitudes of Basin and Range extension can be found to the north and south of the Timpahute transverse zone, but these differences can be accounted for by correlating the thrust segments across the field area (Taylor et al., 2000).

Complex Crater Models

Marine impacts, unlike their subaerial counterparts, involve a water column of varying depth, the presence of saturated unlithified sediments, and often the presence of basement rocks that contribute to a complex morphology (Dypvik and Kalleson, 2010). Detailed analyses of other marine impact sites (e.g., Chesapeake Bay, Mjølnir, Siljan) suggest a specific complex crater morphology (Gudlaugsson, 1993; Kenkmann and Dalwigk, 2000; Dypvik et al., 2004; Poag et al., 2004; Horton et al., 2006). It is assumed here that the Alamo crater has similar complex crater morphology. Terminology used by Poag et al. (2004) to describe features of the Chesapeake Bay complex crater is used here to describe the Alamo crater. This model defines two large-scale features: the transient crater formed directly from the impactor, and the apparent crater formed from the inward collapse of the carbonate platform (Fig. 5).

The apparent crater contains the outer rim fault and the annular trough from its outermost extent inward. The outer rim fault denotes the outermost margin of the apparent crater, signifying the transition from inside to outside the impact crater (Poag et al., 2004). The annular trough is a broad portion of the crater between the outer rim fault and the peak ring (Poag et al., 2004). This portion of the crater forms during the modification stage, as the platform collapses toward the transient crater (Melosh, 1989; Turtle et al., 2005). Often the annular trough makes up >50% of the overall final diameter of marine impact craters (Dypvik and Kalleson, 2010). The boundary between the apparent crater and transient crater is defined by the peak ring and is typically truncated in marine impact craters (Dypvik and Jansa, 2003; Turtle et al., 2005). Inward of this boundary is the inner basin, the deepest excavated portion of a complex crater (Turtle et al., 2005). The center of a complex crater is known as the central uplift and is occupied by uplifted, deep-seated strata deformed during rebound (Turtle et al., 2005).


Four detailed sections were measured from the top of the Alamo Breccia Member to the top of the upper member of the Guilmette Formation (Figs. A1–A6 in the Supplemental File1). In addition, nine previously measured sections were reinterpreted and included in this study (Figs. A7–A17 in the Supplemental File [see footnote 1]) (Anderson, 2008; Thomason, 2010; Myers, 2011). Each measured section log includes the lithology, grain size, matrix, Dunham rock type, sedimentary structures, and fossil content. We interpreted depositional environments from facies and facies associations. Notable sedimentary structures (i.e., desiccation cracks, rip-up clasts) and facies stacking patterns were then used to interpret cyclicity. Depositional sequences and system tracts in this study are identified from facies and cycle stacking patterns, as well as the correlation of mappable sequence stratigraphic surfaces. Sections were correlated along transects using the top of the Alamo Breccia Member as a datum, and tied to sequence stratigraphic terminology established for the Guilmette Formation and underlying Fox Mountain Formation of nearby Egan and Schell Creek Ranges (depositional sequences 1–11 of LaMaskin and Elrick, 1997; Devonian global transgressive-regressive cycles IIa–IIe of Johnson et al., 1996). The Alamo impact occurred during the highstand systems tract (HST) of depositional sequence 3 (HST3), based on an analysis of pre-impact, syn-impact, and post-impact sedimentation across the impacted region using Fischer plots and vertical facies stacking patterns (Rendall, 2013). Transects are roughly parallel and perpendicular (Fig. 1) to the north-trending paleoshoreline (Morrow and Sandberg, 2008). Additional nearby sections (n = 73; Table A1 in the Supplemental File [see footnote 1]) were also considered (Anderson, 2008; Thomason, 2010; Myers, 2011; Retzler, 2013).

We used a regional structural data set to reconstruct modern outcrops to their relative Devonian positions (e.g., Tschanz and Pampeyan, 1970; Bartley et al., 1988; Jayko, 1990; Ackman, 1991; Taylor et al., 1994, 2000; Switzer, 1996; Chamberlain, 1999; Bidgoli, 2005; Bidgoli and Taylor, 2008) using a Python-based geographic information system script developed by Sheffield (2011). This script generates new x-y values for localities within fault-bound polygons according to direction and magnitude of fault offset (Fig. 6) (Tschanz and Pampeyan, 1970; Liggett and Ehrenspeck, 1974; Ekren et al., 1976; Taylor and Bartley, 1992; Taylor et al., 1994, 2000).

Recognition of the outer rim fault is important for making reasonable size estimates of crater diameter (Turtle et al., 2005). We indirectly identified the outer rim fault by facies and thickness patterns of syn-impact and post-impact sediments, providing the means to estimate transient and apparent crater diameter. Direct evidence of syn-impact faulting within the Alamo crater was difficult to discern because of the patchy distribution of localities and the overprint of post-impact tectonic events. Several factors (i.e., impactor angle, velocity, target rock rheology) could have affected the final diameter and symmetry of the Alamo crater, but are unknown and are not considered here.


Depositional Environments

We grouped 18 facies into 7 depositional environments ranging from peritidal to deep subtidal, based on lithology, sedimentary structures, paleontology, field associations, and previous interpretations of similar facies (Table 1). These depositional environments and their stacking patterns are used to infer changes in accommodation along the platform, and to recognize sequence stratigraphic surfaces, systems tracts, and depositional sequences.

Two facies comprise the peritidal environment. Laminated dolomitic mudstone (Fig. 7A) is interpreted to represent a wave-agitated, hypersaline depositional environment with occasional subaerial exposure, such as a tidal-flat setting, given its suite of sedimentary structures. Fenestral lime or dolomitic mudstone (Fig. 7B) is interpreted as peritidal due to the presence of planar laminae, chert nodules, and fenestrae (cf. Mazzullo and Birdwell, 1989).

Two facies compose the shoreface environment. Quartzarenite (Fig. 7C) is interpreted to represent an upper shoreface, marginal-marine setting, indicated by the presence of planar laminations, low-angle cross-bedding, Taenidium burrows, and stratigraphic relationship between peritidal and subtidal facies. Previous studies in the Guilmette Formation interpreted this facies as either peritidal or subtidal (Estes-Jackson, 1996; LaMaskin and Elrick, 1997); however, a shoreface environment is not recognized within those depositional models. Estes-Jackson (1996) described six distinct types of quartzarenite in the Pahranagat Range, distinguished by different sedimentary structures and the presence or absence of burrows. All quartzarenite deposits examined in this study fit within a shoreface environment, eliminating the need for numerous, more specific facies types. Siltstone and/or silty mudstone (Fig. 7D) is interpreted as a lower shoreface, marginal-marine setting based on the presence of Taenidium and Teichichnus burrows, low-angle cross-bedding, planar laminae, and stratigraphic relationship to the quartzarenite facies. LaMaskin and Elrick (1997) described this facies as a tidal-flat setting; a shoreface environment was not included within their depositional model and may explain the discrepancy between depositional interpretations.

Two facies represent the channel environment. Lithic-dominated breccia or conglomerate (Fig. 7E) is interpreted as a high-energy marine bypass channel, supported by a channel morphology, imbricated clasts, and erosional base. Bioclastic-dominated breccia or conglomerate (Fig. 7F) is often found directly atop the lithic-dominated breccia or conglomerate and is also interpreted as a high-energy marine bypass channel. A channel morphology is apparent at two stratigraphic sections that are off the transect lines (HE1 and HE1.5; Table A1 in the Supplemental File [see footnote 1]).

Three facies compose the semirestricted to restricted shallow-subtidal environment. Barren dolomitic mudstone (Fig. 7G) was likely deposited in a saline or evaporitic tidal-flat setting. This interpretation is based on its lack of biota, light color, dolomitization, and association with other peritidal facies. Skeletal mudstone to packstone (Fig. 7H) is interpreted to represent a semirestricted shallow-subtidal setting with increased salinities, given its biota and previous regional studies (see Table 1). Amphipora and stromatoporoid mudstone (Fig. 7I) was deposited within a restricted shallow-subtidal setting, delineated by its preserved laminae, hypersaline biota, and association with other restricted facies.

Two facies define the bioherm environment. Rhodolith grainstone (Fig. 7J) is only found at the Mount Irish reef section (MI1) and was likely deposited as part of the reef complex given its stratigraphic relationship with other bioherm facies, possibly within a back-reef setting (cf. Machel and Hunter, 1994; Ballantine et al., 2000). Stromatoporoid framestone (Fig. 7K) is the most common facies within the bioherm environment and is likely representative of a reef-core setting (cf. Hladil, 1986; Machel and Hunter, 1994). Similar Guilmette Formation facies were interpreted as shallow to intermediate subtidal deposition (LaMaskin and Elrick, 1997; Rendall, 2013); however, a bioherm environment was not distinguished within their depositional models and may account for this difference.

Four facies compose the open, shallow to intermediate subtidal environment. Stromatoporoid boundstone (Fig. 7L) represents an open, shallow-subtidal environment based on the tabular morphology of the stromatoporoids, the presence of open-marine biota, and the interpretations of previous regional studies (see Table 1). Burrowed fossiliferous mudstone to wackestone (Fig. 7M) is sometimes associated with iron-oxidized firmground surfaces and was likely deposited within an open, subtidal setting due to its abundant bioturbation and open marine biota. Siltstone and/or silty mudstone with interbedded mudstone (Fig. 7N) is characterized by alternating beds of yellow to gray siltstone or silty mudstone and dark gray mudstone. Its diverse biota and trace fossils place it within an open, subtidal setting. Other regional studies describe a similarly named facies that differs from our environmental interpretation (Elrick, 1986; Chamberlain and Warme, 1996); however, their facies are devoid of fossils and sometimes include turbidites, unlike the facies in this study. The lower contact of the skeletal packstone to grainstone (Fig. 7O) facies is often an irregular eroded surface infilling preexisting burrow structures. This facies may have been deposited during storm events, given its graded laminae and rare low-angle cross-bedding.

Three facies compose the deep-subtidal environment (Table 1; Figs. 7P, 7R). Tentaculitid silty wackestone (Fig. 7P) is recorded only at one stratigraphic section (MMN4). The tentaculitids are scattered in no preferred orientation, suggesting that they were pelagic and settled out of the water column below storm wave base. Silty-sandy barren mudstone (Fig. 7Q) is interpreted as deep subtidal, given its normal grading, lack of biota, and stratigraphic relationship with other deep-subtidal facies. Mudstone with interbedded siltstone (Fig. 7R) is characterized by thin- to medium-bedded, light to dark gray lime mudstone with very thin interbedded siltstone. The siltstone was likely shed from the shallow platform during episodic storm events or changes in fluvial input, and deposited as thin layers atop the lime mudstone (cf. Elrick et al., 1991).


Both peritidal and subtidal cycles are represented within the two transects. Some successions include massive siliciclastic or carbonate deposits, and appear to be noncyclic. These cases only represent 5% of all successions in this study. Several covered intervals are present at sections MMN4 and MMS2, assumed here as deeper, more recessively weathered units based on the occurrence of barren mudstone float.

Two types of peritidal cycles are recognized. The first type is defined by semirestricted to restricted shallow-subtidal or shallow to intermediate subtidal facies at its base, which shallow upward into peritidal or shoreface facies. The second type is composed solely of peritidal facies and is recognized via changes in sedimentary structures and/or the presence or absence of biota. The base of this cycle type is often a fenestral lime or dolomitic mudstone that is capped by laminated dolomitic mudstone. Cycle tops and bottoms for both types are often marked by desiccation cracks or transgressive lags, respectively.

Subtidal cycles are devoid of peritidal and shoreface facies. Unlike the peritidal cycles, transgressive lags and exposure surfaces are uncommon. Subtidal cyclicity is largely represented by shallowing-upward successions that coincide with coarsening-upward trends. Cycle bases are often marked by shallow to intermediate subtidal and deep-subtidal facies. In some cases, cycle bases are denoted via firmground surfaces that are recognized by an oxidized bedding plane with pre-omission Thalassinoides and Paleophycos burrows. Cycle tops are characterized by semirestricted to restricted shallow-subtidal, bioherm, or shallow to intermediate subtidal facies. Deepening-upward subtidal cycles are less common throughout the transect lines. They generally correspond to fining-upward trends into deep-subtidal facies that are overlain by a coarser, shallow to intermediate subtidal facies.

Sequence Stratigraphy

A compilation of sequence stratigraphic abbreviations, symbols, and facies patterns used to describe the two transects is displayed in Figure 8. The north-south transect (A-A′) and east-west transect (B′-B) are each composed of 7 sections over a reconstructed distance of 55 km and 37 km, respectively (Figs. 9 and 10).

Depositional Sequence 3

Along the north-south transect, HST3 is defined by a series of ∼1-m-thick shallowing-upward cycles (Fig. 9). In the northernmost section (GGS3), these cycles are capped by peritidal or semirestricted to restricted shallow-subtidal facies. HST3 is thicker at GGS3 relative to the southern sections, possibly due to an increase in accommodation space generated during and after the impact. To the south (MMN4 to HE6), HST3 is thinner and shallowing-upward cycles are capped by deep-subtidal and open, shallow to intermediate subtidal facies. Many of the cycle tops were interpreted from firmground surfaces that represent periods of slow sedimentation, and may indicate a more restricted environment; however, at PTN0 in the central part of the transect, a bed of reworked Alamo Breccia Member containing brachiopod and crinoid fragments is present above pristine Alamo Breccia Member deposits. This unit was likely reworked during a storm event and is interpreted to represent an open, shallow to intermediate subtidal environment. As a whole, HST3 is dominated by shallow-subtidal facies in the north and south, while central portions are dominated by deep-subtidal facies (Fig. 11A).

Across the east-west transect, HST3 is characterized by a series of shallowing-upward cycles (Fig. 10). Cycles at the easternmost section (SMFN2) are exclusively peritidal, whereas cycles in the central part of the transect (HCE1 to DMP1) are both peritidal and subtidal. Farther west (MIN2 and MI1), deposits consist of one shallowing-upward subtidal cycle capped by bioherm facies, signifying reef growth along the platform. The westernmost section (MMN4) is composed of two shallowing-upward subtidal cycles with one deepening-upward subtidal cycle between them. Shallowing-upward subtidal cycles are capped by open shallow-intermediate or deep-subtidal facies. Overall, HST3 is dominated by peritidal and shallow-subtidal facies in the east and by deep-subtidal facies in the west (Fig. 11A).

The end of HST3 along both transects is interpreted as a Type 2 sequence boundary–correlative conformity surface (SB-CC-3) inferred from facies and cycles that signify low accommodation space across the platform (Figs. 9 and 10).

Depositional Sequence 4

Along the north-south transect, deepening-upward subtidal cycles, massive noncyclic deposits, and shallowing-upward subtidal cycles are interpreted to represent a platform-wide increase in accommodation space, signaling transgressive systems tract (TST) TST4 (Fig. 9). In the northernmost section (GGS3), TST4 is composed of an ∼30-m-thick, noncyclic shallow-subtidal deposit. To the south (MMN4 and MMS2), deepening-upward subtidal cycles grade into covered intervals believed to represent recessively weathered deep-subtidal facies. TST4 deposits at PNT0 are mostly covered; however, sparse deep-subtidal outcrops are interpreted to mark the maximum flooding surface (MFS-4). By definition, the covered unit between SB-CC-3 and MFS-4 belongs within TST4. Farther south (HN5 and DDB1), deposition of TST4 is indicated by ∼8-m-thick deepening-upward subtidal cycles composed of open, shallow to intermediate subtidal and deep-subtidal facies. At the southernmost section (HE6), TST4 is represented by an ∼10-m-thick subtidal cycle composed of deep-subtidal and open, shallow to intermediate subtidal facies. TST4 deposits thin toward the center of the north-south transect (MMS2 and PTN0) and thicken away from the center. The only exception to this pattern is the TST4 deposit at the southernmost section (HE6), which is noticeably thinner. This may represent sediment starvation of the offshore subtidal setting as an increase in accommodation generation outpaced sedimentation.

TST4 within the east-west transect is defined by a shift into deepening-upward cycles, noncyclic deposits, and subtidal shallowing-upward cycles, similar to those observed in the north-south transect (Fig. 10). The easternmost section (SMFN2) consists of an ∼7-m-thick deepening-upward subtidal cycle of open, shallow to intermediate subtidal facies. At HCE1 and HHN1, the majority of TST4 is represented by thick (>20 m), noncyclic, siliciclastic- and carbonate-rich shoreface deposits. These may represent the formation of a barrier bar along the carbonate platform (Fig. 11B). Farther west at DMP1, TST4 is composed of an ∼12-m-thick deposit of open, shallow to intermediate subtidal and channel facies. MIN2 includes two shallowing-upward subtidal cycles dominated by open, shallow to intermediate subtidal facies. At MI1, TST4 deposits are mainly represented by an ∼25-m-thick noncyclic deposit and overlying ∼5-m-thick deposit, both of the bioherm facies. MMN4 contains several deepening-upward subtidal cycles capped by recessively weathered deep-subtidal facies. TST4 largely records subtidal deposition, with an intermediate- to deep-subtidal setting present at MMN4 (Fig. 11B).

The end of TST4 is marked by MFS-4, interpreted through three observations: (1) the end of noncyclic or deepening-upward dominated cycle deposition, (2) the top of the last shallowing-upward subtidal cycle prior to peritidal cyclicity, and (3) the deepest-water facies recorded at a locality (Figs. 9 and 10). An exception to this is found along the east-west transect at MI1, where MFS-4 has been placed at the top of a rhodolith grainstone deposit (Fig. 10). This deposit is the only recognizable change in lithology within an otherwise noncyclic bioherm interval. Rhodolith-bearing limestones have been interpreted as transgressive marker beds in a study on Cenozoic deposits of New Zealand (Nalin et al., 2008); although many of these rhodolith beds are clast supported, contain numerous other fossil fragments, and/or occur directly atop ravinement surfaces, all of which do not apply here.

HST4 within the north-south transect is indicated by a return to shallow-subtidal cyclicity with minor peritidal conditions, interpreted to represent a loss in accommodation space (Fig. 9). At GGS3, this is recorded in two shallowing-upward subtidal cycles capped by semirestricted to restricted shallow-subtidal facies. Southward, MMN4 deposits record several shallowing-upward subtidal cycles capped by shoreface and open, shallow to intermediate subtidal facies, while MMS2 is composed of shallowing-upward subtidal cycles interpreted from recessively weathered deep-subtidal covered intervals. At PTN0, HST4 deposits are characterized by shallowing-upward subtidal cycles capped by open, shallow to intermediate subtidal facies. Cycle tops at PTN0 are marked by firmground surfaces. HST4 at HN5 and DDB1 is composed of shallowing-upward cycles capped by open, shallow to intermediate subtidal facies. The southernmost section (HE6) records an ∼8-m-thick subtidal cycle capped by open, shallow to intermediate subtidal facies, followed by an ∼30-m-thick peritidal cycle. Overall, HST4 records shallow-subtidal deposition (both restricted and open) in the north and south, and intermediate- to deep-subtidal deposition in central regions (Fig. 11C).

Along the east-west transect, HST4 is marked by the reoccurrence of shallowing-upward peritidal and subtidal cyclicity (Fig. 10). The easternmost sections (SMFN2 to HHN1) are characterized by shallowing-upward cycles capped by peritidal and semirestricted to restricted shallow-subtidal facies. DMP1 is composed of shallowing-upward subtidal cycles dominated by shoreface and semirestricted to restricted shallow-subtidal facies. At MIN2, HST4 is composed of one open, shallow to intermediate subtidal facies. The upper boundary of HST4 is not well constrained here due to the lack of identifiable cycles within this deposit. At MI1, a single bioherm deposit that is above the transgressive rhodolith deposit is interpreted to represent HST4. At MMN4, HST4 consists of shallowing-upward subtidal cycles capped by open, shallow to intermediate subtidal facies. HST4 predominantly records deeper conditions in the west, and shallower conditions in the east (Fig. 11C).

SB-CC-4 is inferred to occur at a shift from carbonate- to siliciclastic-dominated deposition (Figs. 9 and 10). This surface is interpreted as a Type 1 sequence boundary, as indicated by paleokarst filled with terra rossa atop the bioherm deposit at MI1 (east-west transect; Fig. 10).

Depositional Sequence 5

Deposits above SB-CC-4 are characterized by shoreface deposition interpreted to represent lowstand systems tract (LST) LST5 (Figs. 9 and 10). The sequence stratigraphic framework of LaMaskin and Elrick (1997) does not include a lowstand systems tract within depositional sequence 5; however, their locations are in a midshelf position closer to the Late Devonian paleoshoreline (Morrow and Sandberg, 2008). We suggest that LST5 deposits bypassed those midshelf locations and were deposited as a lowstand wedge farther offshore (west). Subaerial unconformities are not recognized in our transects between depositional sequences 4 and 5, except at MI1, suggesting that bypass was submarine (cf. Rendall, 2013). Comparable submarine bypassing across a midshelf environment has been documented in the upper Devonian carbonate platform of western Alberta (Whalen et al., 2000).

Along the north-south transect, the majority of LST5 consists of quartzarenite shoreface deposits (Fig. 9). Finer grained shoreface deposits and channel deposits precede quartzarenite deposition at PTN0, HN5, and DDB1. At GGS3, LST5 is only represented by an ∼1-m-thick silty mudstone deposit. LST5 is primarily represented by a lowstand wedge deposit within the north-south transect (Fig. 11D).

Within the east-west transect, LST5 is recognized by thin shoreface deposits in the east that thicken westward into massive quartzarenite at MMN4 (Fig. 10). The eastern sections (SMFN2 and HCE1) include several shallowing-upward peritidal cycles that we assign to LST5. Below the quartzarenite unit at DMP1 is an ∼5-m-thick channel deposit possibly derived from the bioherm deposits in the west, as indicated by its stromatoporoid- and coral-rich bioclasts. LST5 deposits are not recorded at MI1 because the bioherm was subaerially exposed during this duration, forming a Type 1 sequence boundary (SB-CC-4). Instead, deposition above this boundary consists of several shallowing-upward subtidal cycles capped by bioherm deposits, interpreted to represent TST5.


Seafloor Topography

The north-south transect records a transition from semirestricted to restricted shallow-subtidal, to deep-subtidal, to open, shallow-subtidal facies geographically across depositional sequences 3 and 4 (Fig. 9). In addition, the east-west transect grades from peritidal in the east to deep-subtidal facies in the west geographically across depositional sequences 3 and 4 (Fig. 10). We interpret this pattern to reflect a concave seafloor that is progressively shallower eastward of Tempiute Mountain (Fig. 11). This concavity is consistent with isopach and lithostratigraphic trends in the Alamo Breccia Member seen throughout the study area, and farther to the north and south within the Golden Gate Range and Delamar Mountains (Pinto and Warme, 2008; Sheffield 2011). We interpret this concave pattern as the post-impact legacy of the Alamo crater.

Alamo Crater Features

Along the east-west transect, peritidal and subtidal deposits at SMFN2 transition laterally into an ∼30-m-thick quartzarenite unit at HCE1, coincident with an abrupt thickness change in the Alamo Breccia Member (Warme and Kuehner, 1998; Pinto and Warme, 2008; Sheffield, 2011). We interpret this change in facies and accommodation space as the surficial expression of the outer rim fault (Fig. 12). Localities east of SMFN2 record similar thin Alamo Breccia Member deposits (Kuehner, 1997; Pinto and Warme, 2008).

In the north-south transect, the outer rim fault position has been suggested to exist north of the GGS3 section (Warme and Kuehner, 1998; Pinto and Warme, 2008). At GGS3, depositional sequences 3 and 4 are nearly three times as thick as sections toward the south. We suggest that thicker post-impact deposits here are the result of increased accommodation generated by slip along the outer rim fault to the north (Fig. 13).

Our interpretation of the outer rim fault near the Hiko Hills and Golden Gate Range indicates that the majority of the transect sections are within the annular trough (Figs. 12 and 13). These locations also share similar Alamo Breccia Member lithostratigraphic characteristics (Warme and Kuehner, 1998; Pinto and Warme, 2008; Sheffield, 2011), implying they were formed in a similar portion of the impact crater.

Variations of facies and bathymetry along transects can be explained by faulting within the annular trough, such as between MIN2 and MI1 (Figs. 11 and 12). These variations correspond to thickness changes in the underlying Alamo Breccia Member and lower member of the Guilmette Formation (Figs. 12 and 13) (Sheffield, 2011). Thickness changes in the Alamo Breccia Member may also be the result of one or more of the following: (1) incised valleys formed within the annular trough during resurge events (Dalwigk and Ormö, 2001; Dypvik and Jansa, 2003; Dypvik et al., 2004), (2) differential deposition of the Alamo Breccia Member within the crater, and (3) post-impact structural modification of the crater (Tsikalas and Faleide, 2007).

Currently, no evidence of the peak ring or central uplift has been documented in Nevada. However, the section west of our transects at Tempiute Mountain (TMP in Fig. 1) is the only locality to include evidence of melted and marbleized Alamo Breccia Member (Pinto and Warme, 2008), an indication of proximity to the inner basin (Grieve et al., 1981; Kenkmann et al., 2014). Pinto and Warme (2008) believed this to represent the inner slope of the crater rim that would correspond to the inner basin in the complex crater terminology of Poag et al. (2004). Furthermore, the upper member of the Guilmette Formation at Tempiute Mountain is a mixture of carbonate slope deposits and gravity-flow deposits (Pinto and Warme, 2008) that may have been generated along the transition from the annular trough, across the peak ring, and into the inner basin. If this interpretation is correct, then the peak ring boundary is located between Tempiute Mountain (TMP) and Monte Mountain (MMN4) and would be absent from our transects. The central uplift structure would exist farther west of Tempiute Mountain and coincides with the presumed point of impact near the Quinn Canyon and Reveille Ranges (Morrow et al., 2005; Pinto and Warme, 2008).

Crater Size Estimates

Approach 1. Diameter from Scaling Relationships

Linear scaling relationships have been previously documented between the apparent crater diameter (Da) and transient crater diameter (Dt) in marine impacts (see Fig. 5) (Dypvik and Jansa, 2003; Dypvik and Kalleson, 2010). This relationship, as well as the relationship between the Da and the width of the annular trough (AT), is examined for six well-studied marine impact structures (Table 2). Among these six, Da/Dt values range from 1.9 to 3.0 (mean of 2.5) and Da/AT values range from 3.0 to 4.2 (mean of 3.5). This empirical relationship can be used to estimate the size of the Alamo crater (Da and Dt) based solely on the width of the annular trough (AT).

Data from the six well-studied marine impacts were used to calculate known Da/Dt and Da/AT ratios for marine impacts (Dypvik and Kalleson, 2010). These values were obtained through a stepwise pairing of Da/Dt ratios with each Da/AT ratio, and plotted among actual values (Fig. 14). Values for the Wetumpka crater (Da/Dt ratios < 2.3 and Da/AT ratios > 3.6) were excluded because this impact was unusually small and basement involved, sharing little with the presumed target area of the Alamo impact (Morrow et al., 2005; Dypvik and Kalleson, 2010). From the plot, five calculated (calc) pairs (Table 3) were chosen that represent the highest and lowest possible ratios, as well as the mean (Fig. 14). These pairs were used to estimate the size range of the Alamo crater (Da and Dt) utilizing the AT width and the following two equations: 

The east-west transect in this study only covers ∼37 km (reconstructed) from the outer rim fault to MMN4 within the annular trough. If the boundary of the annular trough (marked by the peak ring) were hypothetically located at MMN4, ATAlamo would equal ∼37 km. This generates an apparent crater diameter (DaAlamo) between 111 and 133 km and a transient crater diameter (DtAlamo) between 37 and 58 km (Table 4). Since the boundary of the annular trough is not apparent at MMN4, these calculations represent the minimum size of the Alamo crater.

Based on the interpretation that the Tempiute Mountain section (TMP) existed within the inner basin of the Alamo crater, the boundary of the annular trough would be between TMP and MMN4. Using the midpoint between TMP and MMN4, ATAlamo would equal ∼55 km. This generates an apparent crater diameter between 165 and 198 km and a transient crater diameter between 55 and 86 km (Table 4).

Approach 2. Diameter of Outer Rim

The apparent diameter of the Alamo crater can also be estimated based on the diameter of the outer rim fault, using interpretations discussed earlier and those made previously (Warme and Kuehner, 1998; Pinto and Warme, 2008). Three control points define the location of the Alamo outer rim fault: (1) between the southern and central Golden Gate Range (GGS and GGC), (2) between HCE1 and SMFN2 in the Hiko Hills Range, and (3) between the Delamar Mountains and the Arrow Canyon Range (ACR) (Fig. 15). A conservative estimate of ∼100 km for the apparent diameter of the Alamo crater was calculated from these points (Fig. 15). As this measurement only captures a chord, not diameter, of the crater, it represents a minimum estimate.

Proposed Crater Diameter and Depth

Morrow et al. (2005) calculated that the final crater could be no greater than 150 km wide based on the lack of crystalline basement clasts within the Alamo Breccia Member and a regional stratigraphic thickness of 6 km of sedimentary strata above basement. This calculation used a mean Da/Dt ratio of 2.5. It is assumed here that the Alamo impact did not involve crystalline basement and a maximum crater size of 150 km is accepted. Therefore, an apparent crater diameter of 111–150 km and a transient crater diameter of 37–65 km are proposed for the Alamo crater (Table 4). This estimate is more than double the Morrow et al. (2005) estimates of 44–65 km for the apparent diameter. Furthermore, this yields an Alamo crater size that is conservatively larger than the Chesapeake Bay marine impact crater (Da = 90 km; Dt = 35 km), one of only eight impact structures discovered on Earth larger than 80 km in apparent diameter (Poag et al., 2004; Dypvik and Kalleson, 2010).

By estimating the diameter of the Alamo crater, the excavation depth (Hexc) can be calculated using the following equation given by Melosh (1989): 

As shown in Table 4, the minimum possible transient crater diameter is 37 km, while the maximum is 65 km. This range yields a minimum and maximum crater excavation depth of 3.7–6.5 km. However, a maximum depth of 6.0 km is assumed, based on available strata atop crystalline basement (Morrow et al., 2005). Consequently, the Alamo bolide would have excavated at least into the lower Cambrian Prospect Mountain Quartzite and possibly into the underlying Neoproterozoic McCoy Creek Group, established by regional stratigraphic thicknesses. This contrasts with excavation estimates by Morrow et al. (2005) of 1.7–2.5 km based on the presence of upper Cambrian microfossil elements within the Alamo Breccia Member. Had the crater excavated into the Prospect Mountain Quartzite or McCoy Creek Group, composed mainly of quartzite and argillite, it would not have left any fossil signatures (cf. Kellogg, 1963; Misch and Hazzard, 1962) within the Alamo Breccia Member. Furthermore, it is unlikely for deeply excavated rocks to be deposited within the annular trough impact breccia (Grieve et al., 1981; Kenkmann et al., 2014), meaning diagnostic lithic fragments containing distinctive or potentially distinctive detrital zircons would not be found within the known Alamo-related deposits. It is possible that these clasts were confined within the inner basin of the transient crater. The best chance of finding such clasts would be within the Tempiute Mountain section or other Alamo-related deposits at or west of Tempiute Mountain. Morrow et al. (2005) reported finding polycrystalline quartz grains in an Alamo-related deposit to the far northwest of Tempiute Mountain near Eureka, Nevada, that may have been derived from the Prospect Mountain Quartzite or even lower.


1. The Alamo impact occurred at the edge of a carbonate platform during highstand deposition; 18 facies of post-Alamo impact deposits record peritidal to subtidal cyclicity in the impacted region.

2. Transects parallel and perpendicular to the paleoshoreline depict a concave seafloor pattern that circumscribes the eastern half of the exposed Alamo impact crater. Significant shifts in accommodation space of impact breccia and post-impact facies reveal the eastern and northern margins of the outer rim fault in the Hiko Hills and Golden Gate Ranges.

3. Identification of the apparent crater margin allows for new size estimations based on scaling relationships observed in other marine complex craters.

4. Conservative estimates indicate a transient crater diameter between 37 and 65 km, and an apparent crater diameter between 111 and 150 km. These values more than double previous estimates based on the biostratigraphy of the impact breccia. If correct, the new estimates rank the Alamo crater as one of the largest Phanerozoic craters on Earth.

5. Bolide impacts have a longstanding influence on sedimentation patterns within a marine setting, shown here across three depositional sequences.

6. The stratigraphic approach of this study can be applied to other marine impact craters to estimate their size, especially those dissected by complex tectonic histories or otherwise lacking seismic data.

This work was made possible by financial support from the National Science Foundation (grant SGP 102484 to Tapanila), the Nevada Petroleum Society (Myers), Idaho State University Graduate Student Research and Scholarship Committee grants (Steenberg and Myers), and Geological Society of America graduate student grant 8819-08 (Johnson). We thank Ben Rendall for help in data collection and interpretation, and Jesse Davenport for field assistance. We also thank Todd LaMaskin, Henning Dypvik, and an anonymous reviewer for helpful comments that greatly improved the clarity of the manuscript.

1Supplemental File. Detailed measured sections (Figures A1–A17) and compiled localities with GPS coordinates (Table A1). If you are viewing the PDF of this paper or reading it offline, please visit http://dx.doi.org/10.1130/GES00964.S1 or the full-text article on www.gsapubs.org to view the Supplemental File.