Human and natural impacts on fluvial and karst depressions of the Maya Lowlands

Human and natural impacts on fluvial and karst depressions of the Maya Lowlands

Geomorphology 101 (2008) 308–331 Contents lists available at ScienceDirect Geomorphology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o ...

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Geomorphology 101 (2008) 308–331

Contents lists available at ScienceDirect

Geomorphology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / g e o m o r p h

Human and natural impacts on fluvial and karst depressions of the Maya Lowlands Timothy Beach a,⁎, Sheryl Luzzadder-Beach b, Nicholas Dunning c, Duncan Cook d a

Georgetown University, United States George Mason University, United States c University of Cincinnati, United States d University of Glasgow, United Kingdom b



Article history: Accepted 13 March 2008 Available online 25 May 2008 Keywords: Erosion Aggradation Maya Lowlands Fluvial Karst Soils

A B S T R A C T This paper begins to differentiate the major drivers and chronology of erosion and aggradation in the fluvial and fluviokarst landscapes of the southern and central Maya Lowlands. We synthesize past research on erosion and aggradation and add new data from water, soils, radiocarbon dating, and archaeology to study the quantity, timing, and causes of aggradation in regional landscape depressions. Geomorphic findings come from many excavations across a landscape gradient from upland valleys, karst sinks, and fans into the coastal plain floodplains and depressions. Findings from water chemistry show that sources in the uplands have low quantities of dissolved ions but water in the coastal plains has high amounts of dissolved ions, often nearly saturated in calcium and sulfate. We found significant geomorphic complexity in the general trends in upland karst sinks. In a few instances, sediments preserve Late Pleistocene paleosols, buried 2–3 m, though many more have distinct middle to late Holocene paleosols, buried 1–2 m, after c. 2300 BP (Maya Early to Late Preclassic). From 2300–1100 BP (Late Preclassic to Classic Periods), the landscape aggraded from five main mechanisms: river flooding, climatic instability, accelerated erosion, ancient Maya landscape manipulation, and gypsum precipitation from a rise in a water table nearly saturated in calcium and sulfate ions. Evidence exists for two or three high magnitude floods, possibly driven by hurricanes. Moreover, lakecore and geophysical studies from the Petén Lakes region have shown high rates of deposition of silicate clays (‘Maya Clays’) starting and peaking during the Maya Preclassic and continuing to be high through the Late Classic. The main driver on upland karst depressions, the Petén lakes, upland valleys, and fans was accelerated soil erosion, but water table rise, probably driven by sea-level rise, was the main driver on the wetlands of the coastal plain because the aggraded sediments here are dominantly composed of gypsum, precipitated from the groundwater. This latter mechanism represents a little recognized mechanism of aggradation over a large region. These large scale environmental changes occurred during periods of intensive ancient Maya land use and climatic instability, both of which may have contributed to erosion by increasing runoff. Despite these geomorphic changes, ancient Maya farmers adapted in several key cases. © 2008 Elsevier B.V. All rights reserved.

1. Introduction Much of the literature on human impacts in geomorphology has focused on relatively modern changes (Hooke, 2000; Ehlen et al., 2006; James and Marcus, 2006), but a diverse and growing literature reveals significant human influence on ancient environments (Butzer, 1992; Denevan, 1992; Redman, 1999; Montgomery, 2007; Beach et al., 2008, in press). The most obvious example is the Mediterranean, where the literature has attempted to differentiate millennia of potential human impacts on geomorphic systems (Butzer, 2005; Beach and Luzzadder-Beach, in press). Parallel to the Mediterranean has been Mesoamerica, which has also experienced millennia of human impacts through its starkly different landscapes, climates, and ecosystems (Beach et al., 2006a). The Maya Lowlands of Mesoamerica ⁎ Corresponding author. E-mail address: bea[email protected] (T. Beach). 0169-555X/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2008.05.019

lies at the heart of the question of human impacts on geomorphology because this region holds such a long history of human land use in its fluvial and karst environments (Beach et al., 2006b). Geoarchaeological research in this region (and other well studied areas like South America and the Mediterranean) is beginning to differentiate the anthropogenic versus natural mechanisms, chronology, and timing of geomorphic change in the context of history and prehistory. Differentiating the mechanisms for fluvial and karst change in the Maya World is a particular challenge, because of the hidden sinks in karst systems, the complicated influence of groundwater, and the lack of published data. Few data exist for water chemistry, sediment load, and the pathways of sediments through these drainage basins. We also have only a broad sketch of the population and land use history for the Precolumbian period, based mainly on spatially limited archaeological surveys that represent only a few Maya sites dating from the ancient Maya period, usually focused on 3000 to 1000 BP. Nonetheless, we do have a growing opus of geoarchaeological survey,

T. Beach et al. / Geomorphology 101 (2008) 308–331

an expanding resource of scholarship on the long history of the region, a concomitant increase in paleoecological and climatological research, a small but growing soils literature, but still very little research on


geomorphology except for karst studies (Day, 2007). The few studies we do have point to increasing evidence for geomorphic change in a variety of sinks in this region, ranging from aggradation behind

Fig. 1. Location map of the Maya Lowlands.


T. Beach et al. / Geomorphology 101 (2008) 308–331

ancient Maya terraces and dams on backslopes and footslopes (Beach et al., 2002), to small upland dolines (Beach and Dunning, 1995), to karst and structural lakes (Wahl et al., 2006; Anselmetti et al., 2007; Dunning and Beach, 2008) to upland valleys, to alluvial fans, floodplains, and to the coastal plain (Pope et al., 1996; LuzzadderBeach and Beach, in press). In this paper, we synthesize and add new findings on soil erosion and aggradation for each of the main geomorphic sinks in the region and provide a new series of data on regional water chemistry, which is one of the major drivers of geomorphic change. Most of our evidence for soil erosion comes from sedimentation studies. We use the characteristics of these aggraded sediments and local and regional water chemistry to differentiate at least five main and two contributing processes of human induced and natural environmental change across this surprisingly diverse landscape. 2. Environments We focus mainly on the Maya Lowlands (Fig. 1) of Mesoamerica, though the Maya Highlands provides an equally interesting and important region because of the role of steep slopes on volcanic and metamorphic bedrock and abrupt climatic and vegetation differences (Dull, 2007). Elsewhere we have described the regional geomorphology of the Maya Lowlands (Dunning et al., 1998), a region that is more complicated in geology, geomorphology, hydrology, and water chemistry than most overviews present. Descriptions often lump the

Maya Lowlands together as one low relief, limestone plateau, weathered into karst hills and depressions with little fluvial activity. Although this description is suitable for part of the Yucatán Peninsula, the southern lowlands in the Tabasco and Campeche states of Mexico, Petén in Guatemala and adjacent Belize, and the river valleys of Honduras and El Salvador contain fluvial systems and landforms (Dunning et al., 1998; Marshall, 2007). The modest amount of geomorphological research, especially fluvial, on this region, however, hinders regional assessment. We present new evidence and review research from the major types of sediment sinks in this region: river valleys around the ancient Maya sites of Copán, Cancuén, Quiriguá, Piedras Negras, Palenque, Xunantunich, and the Three Rivers region of Belize; bajos or karst depressions around the Maya sites of Tikal, La Milpa, Blue Creek, Calakmul, El Mirador, and San Bartolo; and lakes in El Salvador and the central Petén region (Fig. 1). The main foci of this paper are the central and southern Maya Lowlands, from 15 to 19° North latitude and 88 to 92° West longitude and from sea level to about 500 m in elevation (though the Maya Mountains and southern Lowlands have much higher elevations). In the southern reach of this area lie the deep, perennial river valleys of Honduras and southeastern Guatemala that are associated with and controlled by faults associated with the Motagua plate boundary (Fig. 1). The Xibun and Belize rivers in Belize, the Pasión and Petexbatún of Guatemala, the Usumacinta and Candelaria of Guatemala and Mexico, and rivers of northern Belize are fluvial systems that partly drain fluviokarst terrains with relatively high dissolved loads. Parts of these

Fig. 2. Idealized cross-section from upland karst to bajos, river valleys, and coastal plain depressions, showing the location of the Fig. 3 photograph.

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watersheds drain from the central Petén carbonate platform, where some of the studies reviewed here took place on limestone slopes and karst sinks. Several studies have examined the geomorphology of these karst terrains (Siemens, 1978; Day, 1993, 2003, 2007; Reeder et al., 1996; Miller, 1996; Veni, 1996), but this paper focuses on erosion evidence in regional sinks: floodplains, coastal plains, and the seasonally wet ‘bajo’ sinks that cover c. 40% of this landscape (Dunning et al., 2002). These sinks, often 100 m or more above the water table, tend be desiccated during the dry season and drain slowly into seasonal fluvial systems and groundwater in the wet season. The 100–500 m elevation of the karst plateau descends along a series of normal faults northeast and eastward toward the coast, creating a series of escarpments and intervening valleys (Figs. 1, 2). From the central karst uplands river systems initially occupy structural valleys, gathering intermittent upland runoff and perennial groundwaters. From the uplands, rivers descend onto broad floodplains of the coastal plain, gathering other sources of groundwater as they meander north and eastwards to the Caribbean. The Rio Bravo Escarpment (Fig. 3) divides the interior karst uplands from the coastal plain, which spans 100 km across low elevation floodplains, between higher sand ridges, and occasional bedrock outliers. This region lies seasonally under the influence of the Intertropical Convergence Zone and the Bermuda High, which spawn distinctly wet seasons from May to November and dry seasons from December to May. Because of latitude, elevation, and rainshadow, precipitation across the region varies significantly from regions with 5000 mm of annual rainfall in the Highlands and Maya Mountains to those of 500 mm in northwest Yucatán. But most of the region has a much narrower range of about 1200 and 2000 mm of yearly rainfall, and soil moisture regimes are typically Ustic or Udic. All of the soil temperature regimes (except for a small area of higher elevations) are isohyperthermic (mean annual soil temperature is above 22 °C, Van Wambeke, 1987) or isomegathermic (mean soil temperatures above 29 °C). The region also lies sandwiched between two opposite teleconnections of El Niño/Southern Oscillation (ENSO). To the north, El Niño is anomalously wet from October to March, and to the south, anomalously dry from July to March (Ropelewski and Halpert, 1987). These


complications mean the ENSO pattern is variable but significant, as in the extreme aridity and fires of the 1997–1998 El Niño that gave way to torrential rainfall (Golicher et al., 2006). ENSO also influences the tropical storm activity of the region (Boose et al., 2003), with El Niños correlated with decreased magnitudes and frequencies and La Niñas with increases in both (Pielke and Landsea, 1999). Moreover, the Holocene periodicity of ENSO has varied significantly, from little El Niño impact in the early Holocene to an increase after 5–6000 BP (Haug et al., 2001; Koutavas et al., 2002; Magilligan et al., 2008-this volume). These ENSO teleconnections and other teleconnections from the North Atlantic Oscillation drive large interannual rainfall variability. Many factors contribute to ecosystem variability of the Maya lowlands, including soils, the quantity and quality of water, fire regimes, precipitation, and frequency and intensity of tropical storms. Most of the region was covered by tropical forest, including biologically diverse, well drained, high tropical forests, and low and dry, thorny forests. Smaller areas of pine and palmetto savannas also occur in the central Petén and on the sandy coastal plains, interspersed with forested and herbaceous, perennial wetlands (Greller, 2000). Over the last decades settlers have converted much of the area into pasture and other crops (Fig. 3). 2.1. Sea level and water table change In the transition from the Pleistocene to Holocene, fluvial and karst systems changed significantly as sea level climbed 125 m. Blum and Tornqvist (2000) argued that the Mississippi River incised inland 300– 400 km during the Late Glacial Maximum, and the river valleys of the Maya Lowlands would have also incised far inland and to great depths. Ground water tables and river systems then adjusted to these rising Holocene base levels. In the low lying coastal plain, water transgressed far inland, inundating previously active floodplain and cave networks, lifting lighter freshwater tables, and creating a large area of wetlands. Sea levels and water tables continued to rise through the late Holocene at sites at least 40–50 km inland along the Rio Hondo and other Belize coastal plain rivers (Antonie et al., 1982; Bloom et al., 1983; Jacob, 1995; Jacob and Hallmark, 1996; Pohl et al., 1996; Pope

Fig. 3. Aerial photograph of the coastal plain and Rio Bravo escarpment at which the karst uplands starts.


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et al., 1996). Some research from the Gulf of Mexico also suggests 1– 2 m sea level high stands in the mid Holocene (Blum et al., 2002; Otvos, 2001), and Pohl et al. (1996) suggested a possible slight drop in sea level from 5000 to 3000 BP. Studies from northern Belize estimate groundwater tables at about 1 m lower at 3000 BP, about 0.6 m lower by 2000 BP, and about 0.25 m lower by 1000 BP (High, 1975; Pohl et al., 1996; Toscano and Macintyre, 2003; Gischler and Hudson, 2004). Sea level rise also inundated some ancient Maya settlements, including Maya salt production sites by 1 m or more (McKillop, 2005). Some of this submergence of Maya sites is caused by the long-term subsidence of the carbonate platform, which Gischler and Hudson (2004) estimate as 39 to 119 mm ka-1 for the Belize shoreline. An often overlooked part of the physical environment in geomorphology and geoarchaeology is groundwater chemistry (LaFleur, 1984), but in the tropical karst and wetland environments of the Maya Lowlands, we ignore water chemistry at our peril. Luzzadder-Beach and Beach (in press) provide extensive water chemistry analyses to connect sea level, water chemistry, and landscape aggradation in the coastal plain near Blue Creek, Belize, which extends similar, earlier findings (Pohl et al., 1996). 3. Cultural history Cultural history is relevant to geomorphology in a region defined by a cultural stream that has persisted for more than 3000 years. Scholars have divided the pre-Hispanic period into a complex historical sequence (Table 1), which we simplify here to the Archaic (9000– 4200 BP), the Early Preclassic (4200–3000 BP), the Middle Preclassic (3000–2400 BP), the Late Preclassic (2400–1700 BP), the Early Classic (1700–1300 BP), the Late and Terminal Classic (1300–1100 BP), the Post Classic (1100–500 BP), and the Post Conquest period thereafter. Portions of this chronology are artifacts of an outdated literature, particularly the definition of a “Classic” period defined by the erection of stelae with dated hieroglyphic inscriptions, because we now have stelae and inscriptions for the Preclassic and Postclassic. Similarly, early scholars often referred to population levels and cultural accomplishments as having peaked throughout the Maya Lowlands in the Late Classic, followed by population collapse and cultural decline (Turner, 1990). We now consider the course of Maya civilization to have been much more complex. Major population apogees and ensuing collapses occurred in the Terminal Preclassic (1900–1700 BP) and Terminal Classic (1100–1200 BP), with lesser disruptions at the end of the Middle Preclassic and Early Classic. The rise and fall of the Maya, however, was not uniform across the lowlands. Some sites and regions experienced greater growth or more complete abandonment at different times. Some regions were

Table 1 Maya historical periods Years BP (AD 1950)


Erosion Factors

Pre-9000 9000–4200 4200–3000 3000–2400

PaleoIndian Archaic Early Preclassic Middle Preclassic


Late Preclassic

1850–1700 1700–1350 1350–1180

Terminal Preclassic Early Classic Late Classic

1180–1050 1050–700

Terminal Classic Early Postclassic

700–450 450 to present

Late Postclassic Post Conquest, Colonial

Pleistocene/Holocene climate transition Tropical Forest, first agriculture Agriculture, decreased forest Agriculture, decreased forest, Maya clays start to deposit Agriculture, decreased forest, population increase, Maya Clays peak Maya Clays continue Soil Conservation, population increase Soil Conservation, population peak, Maya Clays decrease Soil Conservation, Maya Clays decrease Maya Clays diminish, Reforestation in many places Reforestation in many places Reforestation

abandoned for centuries at a time, whereas others were never abandoned. 4. Soils and soil erosion Small scale, often dated, government soil surveys exist for the five countries we cover, which describe numerous factors of soil formation across these landscapes (e.g., King et al., 1992). One typical sequence includes Rendoll soils across the relatively low gradient backslopes of karst hills and escarpments. These Rendoll catenas often have simple profiles with 20–50 cm of A horizon over C or Ck horizons, over sascab, often a soft calcrete with an indurated or case-hardened surface (Darch, 1981). Variations in the Ck horizon may influence infiltration, percolation, runoff, and erosion, and may represent evidence of past, deeper soil profiles and or Pleistocene aridity. A small number of studies have examined soils across the interior karst plateau region, usually describing relatively thin Mollisols, Alfisols, and Inceptisols on upland slopes and Vertisols or Histosols in depressions and inter grade soils in between (Beach, 1998b; Fernandez et al., 2005; Webb et al., 2007; Johnson et al., 2007a,b). Some Ultisols and Oxisols have formed on older landscapes in the southern Maya Lowlands (Lietzke and Whiteside, 1981; Carlos Donado, 1996; Coultas et al., 1997; Farrell, 2003; Bullard, 2004), but to the south of the Yucatán, where geology becomes much more diverse, soils are too complex to describe in brief. Soil erosion research on the Maya World has not progressed far, despite a long history of scholarly speculation on the relationship of soil erosion and environmental degradation to the ebbs and flows of Maya History and as an explanation for the thin, young Rendoll soils of the region (Beach, 1998b, 2006a,b). Indeed, several early writings ascribed the Maya Collapse of c. 1100 BP to soil erosion or more broadly to environmental degradation (Bennett, 1926; Thompson, 1954; Morley, 1956). Another early hypothesis suggested that degradation by anthropogenic burning formed the Oxisol soils of the central Petén (Stevens, 1964, p. 299). A few researchers have tested these erosion hypotheses with catena studies and some erosion modeling and have found evidence for rapid soil truncation in short time scales (Furley, 1987; Beach and Dunning, 1995; Beach, 1998a; Beach et al., 2006a). Several factors of Rendoll soils make them more erodible in this region. First, the largely clay textures often aggregate into more erodible sizes. Second, the well formed granular structure, which is highly porous, often overlies layers that act like aquitards to infiltration. These layers may be Ck, denser clay horizons, or marly or saprolitic subsoils that are clay-rich and massively structured. Although infiltration starts out high in these soils, even with high tropical rainfall intensity, saturated overland flow and throughflow can perch above these soil aquitards. When other factors of erosion (e.g., deforestation) align, soil losses are very high through sheeting, rilling, and gullying. Indeed, we have observed entire soil profiles erode over several years, though rates of erosion decline as the gullies dissect into the clayey saprolite (sascab) and marl and the Ck horizons cleave off into cobbles and boulders (Beach et al., 2006a). The ancient Maya also had a large range of responses to geomorphic changes in the forms of soil erosion, runoff, and river migration. The Maya started building soil conservation features as early as c. 2500 BP (Preclassic), but most of these systems date to 1700–1100 BP (Classic). They range widely in size, exist in many slope positions, and often contain evidence for intensive cultivation on the soils that have formed behind them (Beach, 1998a; Beach et al., 2002, 2003). Another ancient Maya adaptation was in urban geomorphology, which involved aqueducts and causeways that channelized and diverted water around, away, and under their sites from the Guatemalan Highlands to Copán, Honduras and to Palenque, Mexico. Remarkable stone walled channels at Palenque reflect Maya attempts to counter bank erosion and stream migration (French et al., 2006). Evidence also exists for ancient dams in many places across the Maya

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Lowlands (Siemens et al., 2002). The Maya also built features to capture and funnel water into storage reservoirs (Lucero and Fash, 2006), and we present evidence below for ancient Maya dredging of a reservoir. Although evidence exists for localized Maya dredging, other evidence suggests that sedimentation degraded large bajo depressions and contributed to regional abandonment in the Late Preclassic (Dunning et al., 2002). Accelerated erosion has earned it pejorative connotation, though soil and architectural erosion have deposited beneficial alluvial and colluvial sediments in the footslope and terraces of bajo margins. Indeed, several studies have interpreted the Classic Maya terracing and field wall systems as indicative of intensive agriculture on these formerly degraded zones (Dunning et al., 2002; Beach et al., 2003). Gunn et al. (2002a) study of soil chemistry suggested that Preclassic erosion also delivered more fertile sediments to the massive Bajo Laberinto in southern Campeche, Mexico around the large Maya site of Calakmul, Mexico. 5. Climatic change and erosion Most information on regional climatic change comes from lake and marine sediments, one drainage basin flow study, and speleothembased reconstructions. Studies of lake sediments from the northern Yucatán have provided the main record of climate change in this region based on two main proxies: sediment geochemistry, in the abundance of gypsum (measured as sulfur) to calcite (CaCO3) precipitation, and oxygen isotope ratios (δ18O) in gastropod and ostracod shells. Hodell et al. (1995, 2000, 2001, 2005) have provided important evidence of the general climate tends of northern Yucatán and possibly the broader Maya Region. The cores from Lake Chichancanab (Fig. 1) have provided a proxy record of climate through most of the Holocene, with the following general trends: the lowermost sample is from a terrestrial paleosol, just above this is a sulfur-rich lake with high δ18O ratios from 8200 to 720014C yr BP, followed by low sulfur (S) conditions and δ18O ratios until about 5000 BP, followed by high variability between S and CaCO3 precipitation until c. 2200 BP, followed by two sharp peaks in sulfur in the Late Preclassic and Late Classic that the δ 18O ratios partly parallel in the Preclassic and fully parallel in the Late Classic, and followed by a return to Preclassic levels of δ18O and S and CaCO3 ratios (Brenner et al., 2003). Lake Punta Laguna produced similar results for δ18C ratios for the last 3500 yr, but indicated only broadly drier conditions through the Maya Classic period (highest population densities and land-use intensities as well as the depopulation of the Petén to the south). More broadly, the Haitian site of Lake Miragoane showed a similar sulfur to CaCO3 record, though it included one long period of high sulfur through the Maya Preclassic and Classic. A further study using Gamma ray attenuation bulk density as a proxy for the S proxy revealed Late Preclassic S peaks and two later peaks in drought in the Late Classic (AD 770–850) to Post Classic peaks (AD 920–1100) (Hodell et al., 2005). A new study from the Petén region suggests higher lake levels and precipitation from 2400 to 1800 BP (Preclassic), with lower moisture availability through the Classic period until the present (Rosenmeier et al., in press). Rosenmeier et al. (in press) also show that higher lake levels could have resulted from deforestation or increased rainfall, but the lower lake levels of the Late Classic correlate with strong evidence for the lowest forest cover and increasing aridity. Precisely dated, laminated sediments from the Cariaco Basin, although 2000 km southeast of the Maya Lowlands, provide another marine line of proxy evidence based on variations in sedimentary titanium and iron (Haug et al., 2001, 2003). These sediments are dominated by dark, terrigenous laminae laid down in the wet season and light, organic laminae laid down in the dry season. They interpreted titanium and iron at 2 mm spacing (with a time resolution of 2–4 years) as a proxy of rainfall over the watersheds draining


northern South America. These proxies are low in the Younger Dryas, but rise through the Preboreal to Holocene thermal maximum and decrease through the late Holocene. Three trends particularly stand out in the Late Holocene: high variability from c. 3800 to c. 2000 BP, and stability from 2000 to 1300 BP, low quantities from 1300 to 1000 BP, and the lowest quantities during the Little Ice Age (500– 200 BP). Haug et al. (2001) ascribe the high Ti and Fe in the mid Holocene to high rainfall and runoff, though forest cover would also have been at maximum. They ascribe the 3700 to 2000 BP period of high-amplitude variation to increased ENSO variability. Haug et al. (2003) focus in on the Maya Period and especially the so-called ‘Maya Collapse’, which appears to be a minor event in the data of their early research (Haug et al., 2001) when compared with either the Preclassic instability or the Little Ice Age. Nonetheless, low Ti concentrations, possibly reflecting reduced precipitation, correspond to c. AD 760, 810, 860, and 910, and the Maya Terminal Classic Period. Another approach to understand climatic trends in the Maya Lowlands was linking global insolation, atmospheric patterns, and volcanic emissions with the discharge of the Rio Candelaria, Mexico (Gunn et al., 2002b). The Candelaria tracts these global trends, though the adjacent Champoton follows El Niño patterns. Gunn et al. (2002b) retrodicted these correlations of global temperatures with Candelaria discharges back in time based on solar insolation and volcanic activity. They suggested that forcings correlate with a prolonged period of higher rainfall in the Classic period and dryer conditions in the Terminal Classic. Two paleoclimate records exist for the central Maya Lowlands that may correspond to the Cariaco and northern Yucatán records, the central Petén lakes record and a recent cave speleothem record from adjacent central Belize. For cores from two lakes, Yaeger and Hodell (2008) reported that δ18O declines in the Classic and increases in the Terminal Classic. This implies lower evaporation to precipitation in the Classic and higher evaporation to precipitation (or drought) in the Terminal Classic. But these could have climatic or anthropogenic causes because reforestation also occurred in the Terminal Classic, which, like drought, would possibly increase evapotranspiration and lower lake levels. For a speleothem study from a Belize cave, Webster et al. (2007) used color, luminescence, δ13C, and δ18O as climate proxies. They found climate fluctuation in the Preclassic from drought to pluvial and severe droughts in the Late Preclassic (5 BC and AD 141), then wetter conditions in the Classic Period sandwiching a drought in the Middle Classic (AD 517), and the most severe droughts falling in the Late through Post Classic (AD 780, 910, 1074, 1139). This record may also reveal a drought in the 15th C. AD approximately coincident with a drought in the Mayan books of Chilam Balam of Chumayel and Mani (Folan and Hyde, 1985) and the other core evidence for Little Ice Age droughts. Clearly, more high-resolution paleoclimate studies from this region will help us decipher the spatial and chronological dimensions of climate change and the impact of short-lived, high-magnitude events such as hurricanes. Speleothem studies may also provide more accurate dating correlations between large scale events like hurricanes with episodes of erosion and sedimentation (Frappier et al., 2007). Paleotempestology is new to the Maya world, but one study of historical hurricanes gauged the return period for hurricanes as 1 in 2.4 years and 75 to more than 150 years for Fujita scale 3 events (Saffir–Simpson scale 5 events), wherein most trees are toppled in the affected area (Boose et al., 2003, p. 503). Another recent study on the Belize coast estimated 1 to 1.2 catastrophic storms per 100 years (McCloskey and Keller, in press). The frequency and destructive force of hurricanes led the Maya to perceive these mega storms as the mythological equivalent of chaos (manifest as a writhing, malevolent cosmic serpent) (Dunning and Houston, in press). A more extensive study of Gulf of Mexico sites suggests a return period of 300 years for 4 and 5 Saffir-Simpson scale hurricanes (Liu, 2004). The impacts on geomorphic processes from Hurricane Mitch, especially in Honduras, is an example of how one large scale event could affect the Maya world


T. Beach et al. / Geomorphology 101 (2008) 308–331

Table 2 Soil characteristics of Bajo Majunche, Petén Guatemala Horizon

Dates BP


Depth cm A1 A2 AC C Ab 2C


6.5 6.6 6.4 6 5.5 6.2

Total P

Exch Ca2+

Exch Mg2+

Exch K+

Exch Na+

Organic matter


Clay fine + coarse

mg kg− 1

mg kg− 1

mg kg− 1

mg kg− 1

mg kg− 1

% (LOI)

% (LOI)

% (2 µ − b 1 µ)

153 113 86 64 88 35

8110 7852 7469 6873 7414 6911

1133 979 835 699 748 572

106 88 83 68 71 63

74 79 65 63 71 68

2.8 2 0.8 0.6 7.2 0.4

63 77 77 67 67 66

75 80 79 80 83 77



Table 3 Soil characteristics of Blue Creek: karst doline sink Horizon



Depth cm A, 0–43 Bw, 43–74 C, 74–95 Ab1, 95–130 ACb1, 130–190 2C, 190–224 Ab2, 224–244 3C, 244–290 Ab3, 290–350 ACb3, 350 4C, 350–370

Total P mg kg

Mixed Classic

Mixed Classic

Preclassic 14

C AMS BC 2140–1940


C AMS BC 2475–2195

7.4 7.9 7.9 7.9 8.0 8.0 7.9 8.0 8.0 8.1 8.1


531 – 289 470 – 171 222 – 44 – 39

Exch Ca2+ mg kg 7578 6330 6593 7860 7733 7553 8183 6913 8488 9275 7435


Exch Mg2+ mg kg


449 171 117 144 119 143 163 136 275 348 242

Exch K+ mg kg


1474 433 151 130 125 123 141 104 163 140 132

Exch Na+


Organic matter


Clay fine + coarse

mg kg− 1

CaSO4·2H2O %⁎

% (LOI)

% (LOI)

% (2 µ − b1 µ)

30 40 64 45 36 55 53 65 80 92 184

15 9 9 13 12 12 14 11 14 20 12

11.3 3.8 3.8 4.1 3.9 4.1 3.9 3.7 4.0 4.2 3.5

63 77 77 67 67 66 61 69 56 45 64

29 25 27 36 35 33 39 29 41 52 35



⁎ Crystal-water loss.

climatic instability of the Late Archaic and Maya Preclassic from 3700 to 2000 BP, lower again in the Maya Classic until the Late Classic Maya droughts returned sparser vegetation, lower in the wetter Post Classic, but higher again in the Little Ice Age.

and just how much influence land use can have on high magnitude, low frequency events. Indeed, perhaps paralleling Maya history during the Preclassic, decades of deforestation certainly exacerbated the impacts of Mitch on erosion and mass wasting. Also paralleling Maya history in the Late Classic, the types of conservation the Classic Maya practiced in many places, terracing and silvicultural systems, held 20– 40% more topsoil than adjacent areas without these techniques after Hurricane Mitch's impacts (Krajick, 2001). Devegetation on a large scale has important implications for soil erosion especially when coupled with increased fire in the dry season after hurricanes in this region (Whigham et al., 2003, p. 208). We expect that erosion should increase with drought and hurricanes because of the increased likelihood of fire and its removal of vegetative cover. Several other natural and anthropogenic factors may also affect fire-accelerated erosion, such as the intensity and duration of fire, soil and rainfall characteristics, and land use practices. The relevance of the paleoenvironmental records discussed here to erosion in this region is far from clear because of the inexact connections between climate and erosion, differences in the timescales of these studies, and the extent of applicability of the Cariaco record to the Maya Lowlands, given its distance and the indirect influence of ENSO on the region. If we take these records as possible factors for erosion, soil loss would have been high in the Pleistocene with its sparser vegetation, lower during the full forest conditions of the Holocene Thermal Maximum until c. 3700 BP, higher during the

6. Soil methods From 1991–2007 we sampled and analyzed water and soil chemistry and excavated scores of trenches in transects across the major sinks of the Maya Lowlands, from upland natural karst and anthropogenic depressions to footslopes, to alluvial fans, to floodplains, and to the wetlands of the coastal plain of northern Belize. We synthesize the findings of other regional studies and from our own units and we present new soil data on soil profiles that are representative of these main types of geomorphologic sinks across this landscape (Tables 2–6). The evidence for geomorphic change in each sink includes abrupt shifts in rates of deposition above buried paleosols or increases in sedimentation recorded in sediment cores. In each case, some datable baseline allows us to determine the chronology and quantity of deposition at an area or point. Paleosols provide a useful (but complicated) geomorphic baseline because they represent some measure of stability long enough for pedogenesis to occur and instability in the sedimentation above paleosols. Paleosols, of course, may be unique features that do not represent the history of the

Table 4 Soil characteristics of Blue Creek (Unit 66E): spring escarpment wetland Horizon


Depth cm A, 0–30 Cy, 30–50 Ab/Cy 50–90 2C1, 90–100 2C2, 100–110 Ab2, 135–170

Avail P mg kg

7.8 7.8 7.7 7.8 7.9 7.7

⁎ Crystal-water loss.

15 15 15 18 12 22


Exch Ca2+ mg kg


26580 26360 26800 17632 – 14812

Exch Mg2+ mg kg 630 228 214 208 – 1014


Exch K+ mg kg 21 25 22 28 – 200


Exch Na+ mg kg 83 30 28 27 – 63



Clay fine + coarse


Organic matter

CaSO4·2H2O %⁎

% (LOI)

% (LOI)

% (2 µ − b 1 µ)

65 77 73 9 2 19

3.4 2.1 2.8 1.7 1.3 4.2

28 21 23 83 93 41

27 27 27 15 13 64



T. Beach et al. / Geomorphology 101 (2008) 308–331


Table 5 Soil characteristics of Cacao Creek (Unit BP 1): floodplain wetland Horizon


Depth cm Oa, 0–25 Oe, 25–52 Aby, 52–70 Cy2, 70–75 Cy3, 75–110 Cy4, 110–135 Ab2, 135–160 Cgy, 160–200

6.3 6.3 7.3 7.4 6.0 6.6 7.3 7.6

Total P

Exch Ca2+

Exch Mg2+

Exch K+

Exch Na+


Organic Matter


Clay Fine + coarse

mg kg−1

mg kg−1

mg kg−1

mg kg−1

mg kg−1

CaSO4·2H2O %⁎

% (LOI)

% (LOI)

% (2 µ−b 1 µ)

221 271 85 21 8 64 109 79

13560 21624 24584 25900 27792 22408 15484 16536

1667 1860 2758 2112 1945 2015 2008 2253

43 42 51 33 42 58 75 90

74 70 112 93 75 66 82 102

45 50 50 65 59 48 46 47

18.3 20.5 6.7 3.2 5.7 4.3 5.4 4.6

19 19 14 9 14 13 17 15

78 73 88 68 79 89 88 88



⁎ Crystal–water loss.

broader landscape. Disjunctures in sediment stratigraphy may represent environmental changes as well or, like artifacts or ecofacts, may simply represent a line to date changes in deposition. With these considerations in mind, we chose sites for this review and our field studies that represent broader regional change. For the soil profile studies, we used a range of stratigraphic, geochemical, physical, and chronological methods to identify periods and rates of aggradation in these sinks. Because many of our study sites were perennial wetlands, excavations often required water pumping to collect samples and determine the standard field characteristics: structure, Munsell color, pH, consistence, HCl reaction, magnetic susceptibility, preliminary stratigraphy and horizons (adjusted after micromorphology and laboratory analyses), and further soil descriptions (Soil Survey Staff, 1998). We sampled all soil horizons, soil blocks for micromorphology, all artifacts, and all potential radiocarbon

materials. Field teams used the GF Instruments Magnetic Susceptibility Meter SM-20 to measure magnetic susceptibility at 50 mm intervals on many of the soil profiles. The Milwaukee Soils Lab analyzed samples for the following physical and chemical traits: pH; available and total P, exchangeable K+, Ca+2, Mg+2, and Na+; particle sizes by pipette method; organic carbon and carbonate by loss-on-ignition; and available or total P (Soil Survey Staff, 1996). The Smithsonian Institute SCRME Lab used inductively coupled plasma mass spectrometry (ICP-MS) to determine elemental concentrations on one profile (Cook et al., 2006). The Soils Laboratory at Brigham Young University analyzed some of these soil profiles for carbon isotopic ratios and for elements using inductively coupled plasma atomic emission spectrometry (ICP AES) (Johnson et al., 2007b). The BYU Lab also determined stable carbon isotope ratios (13C/12C) of soil organic carbon after carbonate removal with a Finnigan Delta Plus isotope-ratio mass spectrometer coupled

Table 6 Paleosols and aggradation with depth and dates above paleosols (Ab) or dated charcoal (ch) (updated from Beach et al., 2006a) with new radiocarbon dates shown with beta numbers Location


Pulltrouser and Cobweb, northern Belize Douglas Swamp, Belize Blue Creek Wetlands, Belize Blue Creek Wetlands, Belize

Perennial wetland

Blue Creek, Belize Blue Creek, Cacao Creek, Belize

Upland Valleys Fluvial

Mopan R. Valley, central Belize Rio Bravo, northern Belize Sierra de Agua, Belize Copán, Honduras

Fluvial Fluvial Fluvial Fluvial

Ab, 140 Ab, 120 (80–170) Above flood, 97 Flood, 120 Ab, 135 Ab, 122 Oa, 50 Ab, 80–100 Ab, 150–170 140–190 90–105 74 70–N230

Rio Petexbatún, Petén Nakbe, northern Petén

Fluvial Karst bajo-seasonal wetland

90 50–150

La Milpa Bajo, Belize Calakmul, southern Yucatán, Mexico Bajo La Justa, Petén, Guatenmala Bajo Donato Petén, Guatenmala

Karst bajo-seasonal wetland Karst bajo-seasonal wetland Karst bajo-seasonal wetland Karst bajo-Seasonal wetland

Bajo Mujanche Petén, Guatenmala Far West Bajo, Belize

Karst bajo-seasonal wetland Karst bajo-seasonal wetland

Cancuén, Petén, Guatenmala Cancuén, Petén, Guatenmala

Karst bajo-seasonal wetland Artificial bajo

Bajo Zocotzal Petén Guatenmala

Karst bajo-seasonal wetland

90–170 90–110 80–290 Ab3 = 200 Ab1 = 90 Ab1 = 50 Ab = 90 Ab1 = 30–60 Ab2 = 100–130 Ab2 = 100–130 Ab = 110 ch = 95 ch = 165 Ab1 = 80 Ab2 = 130 Ab3 = 170

Perennial wetland Perennial wetland Perennial wetland

Depth (cm) sedimentation Ab⁎, 140 (75–200)

Dating Paleosol: 14C, cal 2 sigma date

Mechanism 1, 2, 3, 4a


3770–3900 2200–1100 and artifacts 3700–3540 3250–1880 1880–1700 2300–2000 2330–2110 Pre-1700 artifacts 940–760 1170–960 1510–1310 2300–1400 artifacts 2200 –1700 3440–2890 1200 with EPIC modeling; artifacts 2540–1820 1950–2450? and artifacts

2, 3, 1

Pohl et al. (1996) Jacob (1995), Pope et al. (1996) Pope et al. (1996) Luzzadder-Beach and Beach (in press) L-B, B, 2008 Beta 207547 L-B, B, 2008 Luzzadder-Beach and Beach (in press) Beta 207551 Beta 207553 Beta 234154 Holley et al. (2000) Beach et al. (2003) Beach et al. (2003) Wingard (1996:229) Webster (2005: 48) Johnson et al., 2007b Jacob (1995) Hansen et al. (2002) Beach et al. (2003) Gunn et al. (2002a) Dunning et al. (2002) Beta 229813 Beta 232500 Beta 229812 Dunning et al. (2005) Dunning et al. (2002)

3835–3570 2320–2730 1955–1810 12,850–13,090 2790–2710 2100–2020 2770–2850 1710–1570 2380–2200 1420–1270 1250–1320 1300–1360 1830–1560 2200–2015 2850–2700

2, 3, 1 2, 4, 3, 1 2 4 1, 4? 2, 4, 3, 1

1, 4? 1, 4? 2, 3, 1 1 1 1, 3 1, 1, 1, 1,

3 3 4 3

1, 3 1, 4

1 1, 3 4 hurricanes?

Beach et al., 2006b Beta 179794 Beta 179795 Dunning and Houston (in press)

a 1 refers to Maya and/or climate change induced erosion; 2 refers to water table rise of saturated water and evaporate formation; 3 refers to Maya Manipulation or construction; 4 refers to large scale flooding. ⁎ Ab: buried A horizon.


T. Beach et al. / Geomorphology 101 (2008) 308–331

with a Costech Elemental Analyzer. Several archaeological experts identified artifacts, including Drs. Kerry Sagabiel, Lauren Sullivan, Jason Barrett, Jon Lohse, and Collen Hanratty. Beta Analytic assayed all radiocarbon samples using AMS and a few standard radiometric procedures, and we report two sigma (95% probability) dates as calibrated by INTCAL98 Radiocarbon Age Calibration (Stuiver et al., 1998) (Table 6). To counter the problems of dating soils (from bioturbation, old carbon, and carbon mixing), our radiocarbon samples mainly derive from young wood, peat, or charcoal recovered in situ in distinct layers in depositional sequences. We used micromorphology of photomicrographs of soil thin sections to study microstructure and mineralogy. 6.1. Water chemistry methods During the 2006 and 2007 field seasons, we collected 18 water samples at 16 stations across northern Belize. Locations ranged from Chan Chich Creek on the far western border with Guatemala to Four Mile Lagoon on the northeast coast of Belize, and south as far as Hector Creek in the Belize River Valley (Fig. 1). Both field seasons were at the end of the dry season, for temporal consistency with our continuing water quality monitoring program in the region over the last 15 years (Luzzadder-Beach and Beach, in press). We collected our samples in triple-rinsed plastic containers, conducted field measurements of electrical conductivity (EC), total dissolved solids (TDS), pH, salinity, and temperature, then transported the samples to the George Mason University Water Quality Laboratory for analysis. All samples were filtered through 0.45 μm/47 mm Gelman Scientific Metricel membrane filters prior to chemical analysis. Mineral analyses included total hardness (as CaCO3), calcium (Ca+ 2), magnesium (Mg+ 2), sulfate (SO2− 4 ), alkalinity (total as CaCO3), nitrate (NO−3), and sodium (Na+). Table 7 lists the methods and instruments that we used for the water analysis. 7. Results: sediment sinks 7.1. Maya clays The earliest evidence for long-term human impacts on the Maya lowlands came from sedimentary studies of tectonic and karst lakes in the central Petén region of Guatemala (Deevey et al., 1979; Rice, 1996;

Table 7 Laboratory methods for water chemistry Test



Method number, if applicable

EC TDS Salinity NO3–N NO3–N, cont'd Hardness Ca2+ Mg2+ Alkalinity SO2− 4 Cl− − Cl , cont'd Na+ + K+

µS mg L− 1 0/00 mg L− 1 mg L− 1 mg L− 1 mg L− 1 mg L− 1 mg L− 1 mg L− 1 mg L− 1 mg L− 1 mg L− 1

HACH CO150 EC Meter 50150 HACH CO150 EC Meter 50150 HACH CO150 EC Meter 50150 HACH DR850 Colorimeter LaMotte Smart2 Colorimeter LaMotte Direct Read Titrater LaMotte Direct Read Titrater LaMotte Direct Read Titrater LaMotte Direct Read Titrater LaMotte Smart2 Colorimeter LaMotte Direct Read Titrater Calculated from salinity readings Derived from mass balance, K+ is negligible

n/a n/a n/a 8039a 3649-01-SCb 4824-DR-01b 4824-DR-01b 4824-DR-01b 4824-DR-01b 3654-01-SCb 4824-DR-01b 8113b,c n/ac,d


Hach Company, 1998. DR/850 Colorimeter Procedures Manual. Edition 1. Loveland CO: Hach Company USA. b Lamotte Company, 2005. Operator's Manual. Version 2.3. Chestertown MD: Lamotte Company. c Hach Company, 1996. Digital Titrator Model 16900 Lab Manual. Edition 21. Loveland CO: Hach Company USA. d Quality control was ensured by calibrating with standard solutions and testing blank samples prepared with deionized water, and also by calculating a mass balance of cations and anions (Lamotte Company, 2005; Hounslow, 1995; Luzzadder-Beach and Beach, in press).

Rosenmeier et al., 2002; Brenner et al., 2003; Wahl et al., 2006). A typical sequence from these sediment cores included early to middle Holocene organic gyttja with pollen from high forest taxa and low quantities of charcoal followed by declines in forest taxa and the first Zea mays by about 4600 BP, a dense layer of silicate clays dating from about 3000 to 1000 BP (Preclassic through the Classic Maya period) with taxa indicative of diminished high forest, increased disturbance and economic species, and increased phosphorous and charcoal. These sediment sequences then returned to gyttja deposition after 1000 to 500 BP (the Post Classic to modern periods) and also include a return of forest pollen taxa, less phosphorous, and decreased charcoal. A site in the southern margins of the Maya region in El Salvador provides evidence of a similar environmental history but with significant volcanic impacts. Here, near the Ilopango Volcano, Dull (2007) found high rates of sedimentation accompanied with high quantities of Zea mays and disturbance pollen taxa and charcoal in the Preclassic and Early Classic and Late Classic. A hiatus from this trend occurred with (and interestingly before) a large eruption of Ilopango tephra, dated to AD 408–536. The 2000 years of intensive human disturbance (c. 3700 to 1600 cal yr BP) in this region, indicated in multiple lines of evidence, correlates well with the cultural record of anthropogenic disturbance. Human disturbance here also paralleled the Petén records with decline by about 1000 cal yr BP and especially after 500 cal yr BP. Recent research in lake coring coupled with geophysics at the small karst Lake Salpetén in the Petén Lakes Region Guatemala (Figs. 1, 2) is the most detailed and complete lake sediment study thus far to quantify sedimentation chronology through the Maya period (Anselmetti et al., 2007). Because this lake was the main sediment sink in this watershed, the authors could reconstruct the near total flux of eroded soils into the basin through time, and, thus, provide robust estimates of rates of erosion. Based on dating sediment cores from the lake and estimating sediment volumes deposited through time using seismology, Anselmetti et al. (2007) calculated that erosion rose from a background level of 0.25 t ha− 1 to 2.7, 8.9, and 12 through the Preclassic period decreasing to 6 t ha− 1 during the Classic period. These represent counter intuitive rates of erosion because they were highest during the Preclassic when estimated human population densities were low (7–48 people km− 2) and rates of erosion declined during the period of highest estimated population density (up to 248 persons km− 2). The high rates of erosion of the Late Preclassic dropped off by 2.4 times during this time and transition to the Classic Period, but erosion was still 20 times higher than background rates. Other sediment studies from this region, such as Tamarindito in the southern Petén of Guatemala, provide further evidence for Maya Clays and phosphorous enrichment, though not to the extent recorded so dramatically as in the central Petén. Phosphorous concentrations and sedimentation increased 25 times from the Preclassic to Classic periods, and like Lake Salpetén, sedimentation (and by extension, erosion) decreased after about 2000 BP. Also, at Tamarindito charcoal increased significantly in the Late Classic, showing human activity was intensive in spite of lowered sedimentation (Dunning and Beach, 2008). At Laguna Las Pozas, 10 km south of Tamarindito, Johnston et al. (2001) found a brief but significant surge of erosion and sedimentation in the Post Classic around 800–1050 BP. They based their findings on increased in metal elements, magnetic susceptibility, and disturbance pollen. 7.2. Upland seasonal, wetlands: bajos The elevated, interior area of the Yucatán Platform surrounding the central Petén Lakes contains numerous depressions, ranging from small (1–2 km2) karst solution features to large (50–200 km2) structural valleys, all of which are known simply as bajos (Figs. 1 and 2). Drainage is largely internal, but seasonal surface streams, fueled by heavy wet season runoff and perched by dense clay layers, interlink many basins. Today, local variations in hydrologic and

T. Beach et al. / Geomorphology 101 (2008) 308–331

edaphic conditions support a variety of ecosystems including perennial, herbaceous wetlands and various types of palm forests. The most prevalent ecosystem is the tintal bajo: seasonally inundated swamps characterized by Vertisols and stunted, woody vegetation dominated by logwood, or palo de tinto (Haemotoxylum campechianum). Many writers label the “Tintal Bajos” as barren, resource poor environments, although many of the largest and earliest Maya sites in the interior lowlands lie paradoxically along their margins. Several lines of research over the past three decades have attempted to solve this ‘bajo paradox’. Bajos certainly contain clay, seasonal water, and lithic resources, and some scholars in the 1970s used a variety of techniques to suggest that bajos might have been foci for intensive ancient agriculture. For example, previous research has detected ancient Maya wetland field systems and canals in the Bajo Morocoy (Fig. 1) in southern Quintana Roo (e.g., Gliessman et al., 1983). But, further work has demonstrated that canals and field patterns are rare in these elevated bajos, except in this restricted area of Quintana Roo (Pope and Dahlin, 1989; Foster and Turner, 2004; Dunning et al., 2006) and to a much more limited extent in a few other bajos such as the southern margin of the huge Bajo de Azúcar near San Bartolo in the Petén (Dunning et al., in press-a). It is becoming clear, however, that bajos contain a variety of ecosystems and soil associations that are more compatible with traditional Maya agriculture, and that many farmers use them today (Culbert et al., 1989, 1996; Kunen et al., 2000). Indeed, significant ancient Maya manipulation occurred, especially focused on the ecotonal bajo-margins, which are elevated, have accumulating sediments, and hold evidence for water and soil management (Scarborough and Gallopin, 1991; Beach et al., 2002; Dunning et al., 2002; Gunn et al., 2002a). Overall, bajos have surprisingly complex physical, biological, and cultural histories (Dunning et al., 2006). Interdisciplinary research in several bajos in northwestern Belize and northern Guatemala, for instance, have revealed varied environmental histories, at times holding shallow lakes or perennial wetlands that would have held rich resources for the ancient Maya (Castañeda Salguero, 1995; Jacob, 1995; Dunning et al., 1999, 2002, 2006; Beach et al., 2003, 2006b; Hansen et al., 2002). Several examples of bajo geomorphic history serve as examples (Dunning et al., 2005, in press-a). Many bajo sinks experienced ecological and geomorphic changes. One aspect of this is aggradation above buried paleosols (Table 6). Evidence, however, is complicated and variable because of such factors as ancient land use intensity, climatic changes, Vertisol churning of soil profiles, loss of melanization in buried paleosols, and distance, gradients, and flow directions from upland sediment sources. Some of the bajos (Figs. 1, 2) in the upland karst region, such as La Justa and those near the Maya site of Nakbe in the Petén of Guatemala, and Guijarral and the Far West Bajo, near La Milpa in adjacent Belize, provide pollen and carbon isotopic evidence for a shift from perennial wetlands before 2100 BP to drier conditions thereafter as they aggraded from surrounding slope erosion and as the climate became drier (Jacob, 1995; Dunning et al., 2002; Beach et al., 2003; Dunning et al., in press-b). These environmental shifts appear to be synchronous with those experienced along the margins of the Petén lakes that indicate higher lake levels before 2100 BP (Rosenmeier et al., in press). The region around the ancient Maya site of Tikal forms a transition from the karst ridges to bajos. The main source for understanding soils and erosion at Tikal is Olson's survey (1977, p. 16), in which he mapped the soils in the region and estimated 30 to 60 cm (1–2 ft) of coarser sediments buried the pre-Maya, clayey soils of footslopes and depressions. A more recent study has measured c. 1 m sedimentation above paleosols from the Late Preclassic (2500–1700 BP) in parts of the Bajo Laberinto in southern Campeche, Mexico around the major Maya site of Calakmul (Gunn et al., 2002a). In these bajos, buried soils did not have a highly melanized appearance, but the team estimated them by other morphological evidence and an abrupt increase in elemental concentrations using ICP-AES.


In summary, many excavations in the central and southern Maya lowlands display paleosols buried by 0.5 to 2 m, especially adjacent to the surrounding escarpments, where Maya cities and other intensive landuses occurred. Many others examples of early sedimentation exist for the Preclassic such as at the Aguada Catolino, a dissolution doline in the Petexbatún (Dunning and Beach, 2008). Nevertheless, some bajos have little evidence of geomorphic change during the ancient Maya Period, such as the Dumbbell Bajo, parts of the sprawling Bajo de Santa Fe, and larger parts of the Bajo la Justa (Dunning et al., 2006). But those studied thus far represent only a small sample of these large, hydrologically and ecologically complicated karst sinks (Table 6). 7.3. Carbon isotopic ratio examples from depression soil profiles Additional evidence of anthropogenic impacts on erosion comes from soil carbon isotopic studies and an ancient Maya reservoir. Research in the upper Usumacinta River region, near the sites of Piedras Negras and Yaxchilán has examined carbon isotope ratios (δ13C) in soil profiles as evidence for possible ancient Maya use and disturbance (Johnson et al., 2007a). They reasoned that some tropical grasses like maize have a C4 photosynthetic pathway that is both more efficient and less discriminatory toward the heavier 13C isotope than the C3 photosynthetic pathways of forest trees and broadleaf evergreens. Thus soil layers with C4 enrichment reflect inputs from these grasses when these layers were topsoils. Although Johnson et al. (2007a) report few buried paleosols, they did find strong to moderate C4 enrichment in many soils through the top 1–2 m. The horizons of enrichment may coincide with ancient soil surfaces, which may not have been apparent because melanization fades as organic matter decomposes in many soil environments. Two of the bajo soils became more C4 enriched to depths of about 1 m and decreased to become more like a C3 signature below this. A third bajo soil simply increased its C4 signature to the excavations limit at 1.8 m. Other depositional soils also provided similar or even stronger patterns, and one with a buried paleosol at 0.75 m produced the most C4 enriched signature. These isotopic signatures of ancient C4 plant growth occur against a background of earlier C3 forest vegetation, and a subsequent return to forest vegetation after the Maya Terminal Classic, developed in these profiles as a result of rhizodeposition in the root zone, bioturbation, and soil aggradation. Buried soils and the deeply buried C4 carbon isotopic ratios of these depositional soils probably indicate aggradation by c. 1 m by the Maya Classic, but Johnson et al. (2007b) had no dating evidence other than a few sherds from these profiles. We present a new reservoir sedimentation record and five new soil profiles of carbon isotopic ratios from sediment sequences as evidence of the environmental change involved with aggradation (Fig. 4A–C). First, two areas with aguadas (small dissolution dolines) have very similar records despite being located at opposite ends of the Petén with contrasting land covers: recent milpa farming at Cancuén, Guatemala and high tropical forest at Mahogany Ridge, near La Milpa, Belize. We studied two aguadas in low relief areas of the site of Cancuén, an artificial one built in the Maya Late Classic and the second a natural sink. The artificial aguada has infilled to a depth of nearly 2 m since its construction in the early eight century AD. Radiocarbon (Table 6), artifacts (Ohnstad et al., 2004), and geochemistry indicate that 1.3 m of minerogenic sediment was deposited in the ceremonial water feature during at most 110 years of Maya land use, giving a minimal sedimentation rate of 11.8 mm yr− 1. In contrast, postabandonment aggradation during the last c. 1250 years is represented by just 0.70 m of organic-rich sediment, giving a sedimentation rate of 0.56 mm yr− 1. Another aguada at Cancuén was buried by 1.1 m of sediment above an AMS 14C date of 1420–1270 BP on the top of an organic paleosol (Beach et al., 2006b). New carbon isotopic data from this soil sequence (Fig. 4A) is telling: increasing from a mixed tropical forest and recent milpa δ13C signature (−25.05) at the surface to a signature that shows considerable input of C4 species in the Late


T. Beach et al. / Geomorphology 101 (2008) 308–331

Fig. 4. Carbon isotopic ratios (δ13C) of four depression soil profiles. A. Upper left, soil profile from a karst sink at Cancuen; B. upper right, soil profile from Mahogany Ridge, Belize; C. lower left, three soils profiles from near the site of La Milpa, Belize, D. lower right, soil profile from the Cacao Creek, Belize floodplain in the coastal plain.

Classic period (−17.75), but then a significant decline (−25.83) in the earlier subsoils. Thus, the bulk of sedimentation at Cancuén occurred when the landscape had a much larger presence of C4 species during the Maya Late Classic. These examples from Cancuén aguadas indicate a minimum decline from the Late Classic to Post Classic in the sedimentation rate by 21 times, which is a comparable magnitude to Salpetén's 24 times decrease over the same time period. Aguadas from near the site of La Milpa, Belize provide contrasting histories. The first at Mahogany Ridge (Fig. 4B) had a similar record to Cancuén, with a paleosol buried by 0.92 m in the Classical period (dated by artifacts) and a steady increase in the δ13C from the modern topsoil to the buried soil, increasing from a typical tropical forest signature (−27.33) to a signature that shows high input of C4 species in the Classic period (−21.67), but then a significant decline (increase in C3 species organic matter) in the earlier subsoils (−24.61 at Mahogany Ridge). Apparently, the Cancuén and Mahogany Ridge soils experienced Classic period aggradation, during which time significantly more C4 vegetation (such as Maize) was contributing to the soil organic matter. Less apparent were three other depression profiles around La Milpa (Figs. 1, 4C). Here two paleosols buried at depths of about 1 m and dating to the Preclasssic at the La Milpa Aguada (3180–

2980 BP) and bajo (3695–3400 BP) display increases in the δ13C from the topsoils to the Preclassic paleosols by 3.23 and 3.54 respectively, which is a magnitude within the range of natural fractionation down soil profiles (Johnson et al., 2007b). In both cases, nonetheless, δ13C decreased again into the subsoils to nearly the topsoil signature. The least compelling was a profile with a Classic paleosol (radiocarbon dated to 1585–1295 BP) buried 0.62 m behind an ancient Maya terrace. The δ13C increased by only about 2 ppt, a quantity well within the range of natural fractionation and indicative of continuous C3 plant inputs over time. 7.4. Geoarchaeological example of the Bajo Donato The Bajo Donato is a small (c. 2 km2) depression perched in the rolling uplands seven km northeast from the site of San Bartolo, Guatemala (Fig. 1; Table 6). The site of San Bartolo includes monumental architecture dating to at least 2800 BP and the earliest known examples of Maya writing and mural painting, dating to the first century BC (Saturno et al., 2006). Bajo Donato has surface drainage via a shallow arroyo into the Bajo de Azucar to the north. Modern ecosystems include palo de tinto, low scrub forest vegetation,

T. Beach et al. / Geomorphology 101 (2008) 308–331

which is flanked by mixed palm forest or an abrupt transition to upland forest on karst ridges. Ancient Maya terraces cover the southern and eastern slopes above the bajo. The Aguada Tintal (a small, human-modified pond) lies in the northwest portion of the Bajo Donato. A sediment core extracted from the aguada in 2005 found agricultural pollen (Zea mays, manioc (Manihot esculenta), and cotton (Gossypium hirsutum) and evidence of forest clearance in sediment with a calibrated age range of 2730– 2360 BP (Dunning et al., 2005). Sediment from about 2100 to 1200 BP was absent and perhaps removed by ancient Maya dredging. Excavations in 2007 revealed that the Maya had begun deepening the aguada and constructing a stone-lined berm on its margins in the Middle Preclassic, burying an adjacent soil surface that produced a calibrated bulk humate radiocarbon date of 2840–2730 BP (Dunning et al., in press-a). Nearby low, rubble berms and modified surfaces may have been used for local drainage modification and agriculture. Another excavation (Fig. 5A) in 2007 near the southern end of the bajo revealed multiple episodes of pedogenesis and aggradation (Dunning et al., in press-a). At a depth of nearly 2 m lies a highly compressed organic Ab2 (or histic) horizon that produced a calibrated radiocarbon range of 12,850–13,090 BP. Although this buried material did not preserve pollen or phytoliths, it did retain some ephemeral impressions of leaves and grasses and organic matter here produced a δ13C value of −24.8 (a middle range carbon isotopic ratio indicating


either a mix of woody and herbaceous vegetation or seasonal wetland (Johnson et al., 2007b). We interpret this to be a relict organic surface buried rapidly (to preserve leaf imprints), then further buried by continued aggradation later in the Holocene. Cowgill and Hutchinson (1963) reported a similar Late Pleistocene (11,500–11,860 BP) black horizon at a depth of 3.5 m in the Bajo de Santa Fe near Tikal. More representative of bajo profiles was another highly compressed organic layer at a depth of 0.9 m in the small Bajo Majunche, southwest of San Bartolo in 2005 (Table 2; Fig. 6). Bulk organic matter sampled from here dated to 2770–2850 BP, and preserved pollen indicative of a shallow lake or perennial wetland at that time (Dunning et al., 2005). Analogous buried histic deposits with perennial wetland pollen have also been found in the Bajo de la Justa in northern Guatemala and the Far West Bajo in northwestern Belize, but dating to 1800–2000 BP (Dunning et al., 2002). In contrast, the buried Pleistocene horizons in the Bajo Donato and Bajo de Santa Fe were unlikely from perennial wetlands, since these Late Pleistocene environments were much more arid (Leyden, 1984; Brenner et al., 2002). Although bajos have notoriously variable hydrologies, at least some were also well drained, judging by the presence of a buried late Pleistocene Oxisol in the Dumbbell Bajo in northwestern Belize (Dunning et al., 2006). Lying above the Ab2 horizon in the Bajo Donato was clayey mineral sediment and another buried organic rich paleosol. This paleosol is

Fig. 5. Illustrations of three Bajo soil profiles from the Petén, Guatemala. A. Bajo Donato 1 shows a Vertisol with two buried paleosols. B. Bajo Donato 2 shows a Rendoll with one buried paleosol (redrafted from Dunning et al., in press-a). C. Bajo Zocotzal near Tikal, Petén, Guatemala, illustrating two high energy events and three paleosols (Re-drafted from Dunning and Houston, in press).


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Fig. 6. Bajo Majanche soil profile and photograph, Petén, Guatemala (redrafted from Dunning et al. in press-b).

topped by the Ab1 horizon, with a calibrated radiocarbon range of 2950–2750 BP. This surface was in turn buried again by clayey mineral sediments of the upper horizons. Subsequent hydrologic conditions within the bajo and the uneven distribution of accumulating masses of sediment combined to induce substantial argilloturbation, clearly seen in the breakage of horizons, heaving, and slickensides developed throughout the soil profile (Fig. 5). Several studies have documented Vertisol development in aggraded clay sediments of many bajos in the Maya Lowlands (Jacob, 1995; Dunning et al., 2002, in press-b; Beach et al., 2003, 2006a). Further excavation of the base of the ridge forming the southern margin of the Bajo Donato exposed another Ab horizon at c. 0.5 m depth (Fig. 5B; Donato 2). Bulk carbon from the upper Ab produced a calibrated radiocarbon range from 2150–2250 BP (Table 6). Colluvial deposition with numerous, coarse limestone clasts probably buried this horizon from upslope erosion after around 2000 BP (later Late Preclassic) (Dunning et al., in press-a). The ridge upslope from this unit has a slope of 5–8°, and the Maya modified the slope with a series of contour agricultural terraces. Excavations of these terraces revealed a single low wall of stacked stones backed by a wide berm of rubble. A carbonized maize kernel, recovered from c. 0.6 m behind one of the terraces, produced a calibrated radiocarbon range of 1860–2000 BP (Table 6), similar in age to another Ab horizon (1930–2130 BP) behind a buried ancient Maya terrace in another Arroyo at Xultun, 8 km south of San Bartolo. These lines of evidence suggest that the terrace construction was also sometime around 2000 BP, possibly connected to slope erosion (Dunning et al., in press-a). The data from the bajo, aguada, and southern flank suggest the following sequence of development. At the end of the Pleistocene the Bajo Donato had a mixed forest vegetation or was a wetland with a thick O horizon. During the early to mid-Holocene, the depression aggraded rapidly with clayey sediments, followed by a period of geomorphic stability. By 3000 BP, the bajo included mixed forest and wetland vegetation. Shortly thereafter, Maya settlers arrived in the area and began clearing forest — probably for swidden cultivation. The Maya began mining a chert-rich stratum exposed near the surface of

the Bajo Donato during the Middle Preclassic (c. 2700 BP) and after 2600 BP they adapted the former quarry to enhance its waterretention capacity, thus creating the Aguada Tintal. Middle and Late Preclassic farmers in the area were cultivating maize, manioc, and cotton among other crops. Accelerated erosion ensued on the surrounding uplands inducing renewed aggradation in the bajo. By the first century AD, Maya farmers sought to stabilize soils on the southern flank of the bajo using agricultural terracing. The area (including the center of San Bartolo) was apparently abandoned around 1850 BP for unknown reasons, but was reoccupied by a smaller agricultural population in the Late Classic (c. 1250 BP) for over a century, during which time dredging occurred, perhaps to enhance its water-holding capacity. Several anthropogenic factors drove bajo aggradation in the interior southern and central lowlands including urban construction, deforestation for lime mining, agricultural erosion, and ancient causeway construction. In general bajos aggraded from 1–2 m between the Preclassic through Classic Maya times near Tikal and Mirador (Beach et al., 2003, 2006b). But erosion here was also natural (perhaps accelerated by humans) such as the middle Holocene aggradation of paleosols and two high energy deposits (high magnitude, low frequency events). In the Maya Lowlands, at least during the late Holocene, high magnitude events are most likely to arrive in the form of tropical storms, which are common in the region and the ancient Maya recognized in their writing (Dunning and Houston, in press). Two separate deposits in different bajos (the Far West Bajo near La Milpa, Belize, and the Bajo Zocotzal near Tikal (Fig. 5C) stand out because of the thick, high energy (coarse grained) facies within mostly fine grained facies. These high energy deposits account for about half of all Late Holocene sediments in these excavations. The example at Bajo Zocotzal displayed two high energy strata, 0.8–0.9 m thick, bracketed with three paleosols formed in low energy clays. The upper paleosol dated to the Early Classic (1830–1560 BP), the middle paleosol that separates the two high energy deposits dated to the Late Preclassic (2200–2015 BP), and the lowest paleosol dated to the Early Preclassic (2850–2700 BP) (Dunning and Houston, in press).

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7.5. Fluvial environments Along its peripheries, the Maya Lowlands has many fluvial and fluviokarst watersheds such as the Pasión, Belize River, Candelaria, Motagua, Polochic, and Usumacinta (Fig. 1), but scant fluvial geomorphological research exists for these watersheds. Most regional river systems have low gradients, wide seasonal flow variability, and carry high dissolved loads. Flows tend to be highest from June to December, when they often back up and reverse stream directions. Dry season flows are from base flow or from high elevation runoff and, thus, stream loads are highly seasonal. Archaeological reports from the region provide some evidence of past fluvial activity. For example, early research on the Belize River found that Preclassic structures in the upper river terraces near the site of Barton Ramie were built into or on top of a 0.70 m thick paleosol, which in turn was buried by c. 1 m of brown, and then black, clay-rich topsoil (Willey et al., 1965). Furthermore, subsequent structures from the Classic period are founded within the upper sediments; thus, sedimentation occurred in the Preclassic and in or after the Classic period. Willey et al. (1965) ascribed this aggradation of the upper terrace (which previously could not be reached by floods) to increased flood levels caused by accelerated runoff from watershed deforestation (Olson, 1981). They also noted a tendency for sites to be farther uphill after the Preclassic. Holley et al. (2000) also found evidence for preclassic alluviation near the Maya site of Xunantunich with an electromagnetic conductivity survey and six excavations. Generally, they found well formed topsoils and 1.4 to 1.9 m of high energy sediments (cobble, gravel, and sand) burying clayey palesols. They also used artifacts to date the paleosol and the aggradation to the Preclassic and ascribed it to large-scale, human induced flooding in the watershed (Holley et al., 2000, p.122). These high energy events could have been one or multiple events, but these sediments provided no evidence for periods of stability. One study of the Xibun River in Belize found buried soils along two of its Holocene terraces, identifying an upper terrace paleosol buried by c.1 m of alluvium sometime in the middle Holocene and a lower terrace paleosol buried by about 0.60 m of alluvium. These terraces have truncated paleosols with Bt horizons covered by 0.60 m of sand and silt with weak pedogenesis. This deposit also buried Late Classic architecture (1120–1000 BP) that is 5–6 m above the modern surface of the river, which dates this sedimentation to sometime after abandonment, c. 1000 BP (Bullard, 2004, p. 319). Lietzke and Whiteside (1981) report a similar but undated Ab sequence buried 69–86 cm from the Swazi River floodplain in southern Belize. The footslopes and river valleys around the Maya site of Copán provide an example of long-term sediment sinks. Though no study has expressly investigated the sediment budget of these valleys, Turner et al. (1983) did study the Holocene history of the river. Their investigations indicate that remnant Late Pleistocene valley fills are preserved in a set of upper alluvial terraces, carved by early to midHolocene entrenchment, upon which have developed Oxic Ustropepts. A more extensive, lower terrace represents the pre-Maya valley floor. Gully exposures and pits in this terrace often revealed two buried soils: an upper paleosol that was buried by historic or modern alluviation, and a lower paleosol (typically at a depth of 0.8–1 m), which is the actual pre-Maya or Preclassic soil (a Mollic Ustifluvent), artifact-dated to c. 3000–2000 BP (see Turner et al., 1983, Fig. T-23). This surface was variously buried and destroyed by frequent flooding, channel migration, and floodplain sedimentation in the period from 3000 to 1500 BP, probably attributable to land clearance and agriculture as population expanded within the valley. The Copán Maya responded to the increasingly erratic river by diverting it sometime after 1300 BP, though the system apparently fell into disrepair after the site declined after 1200 BP. From 1000 to 500 BP, the Copán River started to incise into the previously aggraded floodplain and formed the second lower terrace. Eventually, the


river also began incising into the Copán Acropolis, necessitating the diversion of the river again by archaeologists of the Carnegie Institution of Washington working at the site in the early 20th century. Wingard (1996) also studied soil sequences around Copán, focusing mainly on the top 1 m. One of his soils pits unveiled a clay and clay loam-textured paleosol buried by 0.7 m of sandy loam sediments in the Copán Pocket. Olson (1981, pp.113–114) also reported a buried A soil horizon in alluvium with Maya artifacts at 1.07 m in the nearby Rio Amarillo and at the Valle de Naco region in Honduras. Moreover, Wingard (1996) concluded that some Late Classic house mound groups were covered by at least 2 m of sediment in footslopes and floodplains around the main site of Copán, which Webster (2005, p. 48) provides in a diagram. Both Wingard and Webster suggest that accelerated erosion and deposition resulted from the expansion of Maya agriculture onto surrounding hills and the cultivation of steeply sloping Rendolls and various types of shallow Inceptisols. This general scenario of Late Classic accelerated erosion and sedimentation is also supported by a coring study of the Aguada Petapilla, a pond within a tributary valley (Abrams and Rue, 1988). The main research on the Motagua River (Fig. 1), Guatemala has focused around the ancient Maya site of Quiriguá (Ashmore, 2007). A modern irrigation project exhumed several 2 m deep canals across the floodplain here, which provided windows into regional archaeology and geomorphology. The floodplain revealed that 1–2 m or more of alluvium enveloped all but the major architecture of the site (built from c. 1550 to 1200 BP). Ashmore (2007, p. 23) estimated the highest rates of alluviation were 0.22–0.29 mm yr− 1 from 1400 to 1100 BP. Estimated alluviation declined after this to 0.05–0.06 mm yr− 1 at greater distances away from the river channel and to 0.18–0.2 mm yr− 1 near the channel. Excavations did not extend below to earlier levels. Soil studies across the Rio Petexbatún (Fig. 1) have found some evidence for aggradation associated with Maya land use in karst sinks, floodplains, and footslopes (Beach, 1998a; Dunning and Beach, 2008). Sediment profiles of the Rio Petexbatún floodplain include buried anthropogenic soils at depths of 1.1 and 0.9 m. The latter paleosol near Aguateca, dated to 2540 to 1820 BP, demonstrated strong C4 carbon enrichment down the profile to the paleosol (Johnson et al., 2007b). Scatena and Golden (2003) studied the general geomorphology and environment of the Usumacinta River in the northwestern Peten. Initial geoarchaeological investigations around La Joyanca along the Rio San Pedro Matir, a major tributary of the Usumacinta in the northwestern Petén (Fig. 1), have uncovered evidence for dynamic environmental changes associated with the Pleistocene–Holocene transition and with the ancient Maya occupation of the region (Métailié et al., 2003). In this region of muted ridge and valley terrain, late Pleistocene stream channel draining ridges were aggraded, probably as the result of base-level changes in the Holocene. In a lake core study, Carozzal et al. (2007) found evidence for Zea mays agriculture at the bottom of their core, dated to 2600 BP, evidence of sporadic erosion in increased magnetic susceptibility and clay deposition spikes in the Preclassic, decreased evidence in the Late Preclassic, but the highest erosion rates in the Early Classic. They also found evidence for increased erosion in the Post Classic about 800 BP in colluvial and lacustrine deposition. Hence, the handful of studies that have assessed fluvial geomorphology in the Maya Lowlands have found evidence for erosion and sedimentation coinciding with Preclassic and Classic periods, but fluvial research is in its infancy in this region and often centered around major archaeological sites, which may make this research non representative. 7.6. Three rivers region and the coastal plain wetlands We consider the Three Rivers Region around the sites of Blue Creek, Chan Cahal, and La Milpa (Figs. 1, 3) as a microcosm of the southern and central Maya Lowlands. It forms an ecotonal transition


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from the karst upland ridges and valleys of the interior, the confluence of three main river valleys, and the confluence of groundwater aquifers at the last major escarpment between the mostly karst interior uplands and the fluvial and karst coastal plain. This region has allowed us to study sediment sources alongside water chemistry and compare these with the depositional record in a series of sediment sinks from upland depressions, alluvial fans, floodplains, and perched coastal plain depressions (Luzzadder-Beach and Beach, in press). The bajos discussed earlier extend to near the edge of the Rio Bravo escarpment (Fig. 3). Along the escarpment are karst ridges, dry karst valleys, dolines, and alluvial fans. Excavations and exposures across upland karst valleys near Blue Creek display rapid, recent gully erosion that reaches up the backslopes. One modern gully formed since European settlement in 1958 has incised by as much as 3 m in depth and over 125 m in length through the sediments of a non-channeled valley (Luzzadder-Beach and Beach, in press). The gully exposes a deep, modern Rendoll topsoil (0.2–0.3 m thick, black, and granular structured) and an even thicker and darker paleosol buried at 1.22 m, which sandwich high energy colluvium and alluvium (Fig. 7, photo). The sequence was dated with artifacts: Late Classic and mixed ceramics in the upper 0.98 m and strictly Late Preclassic ceramics through the next 0.24 m of more melanized colluvium into the upper paleosol (Fig. 8A). Luzzadder-Beach and Beach (in press: e.g., see unit 66W) demonstrated that some alluvial fans at the contact between the escarpment of the upland karst region and the coastal plain produced a sequence that parallels this karst valley because they also had alluvial and colluvial sediments sandwiched by a well formed topsoil and a paleosol dated to the Preclassic by artifacts and radiocarbon dates (Fig. 8B). Webb et al. (2007) described similar aggradation over two footslope paleosols in northern Guatemala, although without chronology. These soils also have significant delta carbon enrichment down to the paleosols. Another fan (66S), however, supplied a more complex picture with a series of paleosols from 3.1 to 1.47 m dated to the Archaic and Preclassic and the sediments above 1.4 m dating only to the Classic and later periods (Fig. 8C). Artifacts and radiocarbon dates show surprising fan instability before the Classic Period, but also Classic instability capped by cumulic A horizons. An upland doline excavation at 66M (Table 3) provided a similarly complicated environmental history to the alluvial fan at 66S (Beach et al., 2006a). This nearly 4 m deep unit had a paleosol near the bottom at 3.5 m that yielded an AMS date of 4425–4145 BP on charcoal. Charcoal from a paleosol at 2.9 m

depth yielded an AMS date of 4090–3890 BP, and Preclassic artifacts dominate the sequence below a paleosol at 1.30 m, which in turn lies below alluvial and colluvial sediments that have mixed, dominantly Classic period artifacts through the well formed, cumulic topsoil. Hence, this complex sink had high deposition in the Archaic and Preclassic periods, and further deposition largely in the Classic period judging by the ceramic chronology and the degree of topsoil formation. The coastal plain wetlands (c. 7–12 masl) are another geomorphic sink that provides clear evidence of widespread aggradation (Luzzadder-Beach and Beach, in press). These wetlands and rivers flow toward the Rio Hondo and on to Chetumal Bay, 80 km from Blue Creek with a regional landscape gradient of about 0.1 mm km− 1. We have excavated many soil units through these wetlands to understand the chronology and processes of formation. Unit 66E was a representative excavation (10 m long, 2 m deep) of the soils, sediments, and timing of aggradation in this area (Table 4; Fig. 8D). As with all other units in these depressions, Unit 66E preserved the following profile (Fig. 8D): a thick, well developed paleosol (Eklu'um Paleosol) buried by 1.3 m of sediment. The paleosol had a thick, clayey Abss horizon topped in places by a 0.1–0.2 m thick O horizon and topped here and in all of our excavations by 0.09 m of carbonate sand. This sequence is in turn buried by about 1 m of sediment that is generally about 75% gypsum particles (clay and silt sized). The soil profiles of the upland sites (Figs. 5–7, 8A) look similar to these coastal plain sites (Figs. 8D–F, 9), but the uplands are aggraded by dominantly carbonate rich sediments and the coastal plains are aggraded by dominantly gypsum rich sediments. Four additional lines of evidence differentiate the formation of these coastal plain sequences: the profiles of magnetic susceptibility (Fig. 10 A, C), general soil characteristics, elemental analysis of the soil horizons (Fig. 10B), and micromorphology. In all excavations, magnetic susceptibility is relatively high in the modern topsoils, decreases through gypsic and carbonate sediments that overlie the paleosol, and increase by a factor of 4 or 5 through the paleosol (Table 4). In all excavations the surface soils are clay loams but the gypsum and carbonate layers are coarser in texture, often loam or silt loams. Elemental analysis by ICP-MS from an excavation adjacent to 66E provides further evidence. Here we used the same techniques as discussed in Cook et al. (2006) and display elements logarithmically to present them on one graph (Fig. 10B). The content of iron, for example, tracks the magnetic susceptibility patterns, with a three fold increase in iron in the topsoil and Eklu'um

Fig. 7. Photograph of gully, showing soil profile with paleosols, northern Belize uplands.

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Fig. 8. Generalized buried soil sequences across the Three Rivers Region, northern Belize. 323


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Fig. 9. Photograph of excavation unit 66E, showing soil profile with paleosols, northern Belize coastal plain wetland.

Paleosol over the gypsic and carbonate layers (Fig. 10B). The same pattern holds for most elements, such as zinc and lead, though barium and rubidium also increase, along with magnetic susceptibility, by

about two times through the transient Ab horizon at about 0.7 m. Strontium is the only element that has an opposite pattern through the profile: 3.5 times higher in the topsoil and gypsic layers than the paleosol. Micromorphology of these horizons corroborates the soil chemistry: clays and carbonates dominate the paleosol, carbonate also dominates the flood horizons, gypsum comes to dominate the upper 1 m, but calcite increases into the modern topsoil (Table 4). The lower sequence of the paleosol and the carbonate sands represented a carbonate and clay rich environment that was buried by gypsum-rich sediments. Luzzadder-Beach and Beach (in press) present numerous radiocarbon dates from similar soil sequences revealing that the Eklu'um Paleosol is Preclassic in age, ranging from 3250 BP to 2110 BP through the Abss (Figs. 8, 9). Three lines of evidence indicate the Abss paleosol was stable for a long period of time: vertical thickness, high clay and metal elemental content, elevated magnetic susceptibility, slickenside formation, and the soils preservation of archaeological evidence of human activity (copious charcoal, economic pollen, and artifacts). With a new AMS date and other previously published dates from an adjacent trench with the same stratigraphy, charcoal from the Ab horizon just beneath the carbonate sand layer dates to 2330 to 2110 BP, charcoal from the sand layer dates to 2300–2000 BP, and charcoal just above the sand layer in the gypsum rich layer dates to 1880 to 1700 BP (Fig. 8E). Hence, the paleosol became buried by the sand layer sometime from 2300 to 2000 BP at the 95% probability (Late Preclassic). This sand layer represents the highest energy event in the last three millennia and makes up about 7% of total aggradation (Luzzadder-Beach and Beach, in press). The upper gypsum rich horizons started to form between 1880 to 1700 BP (latest Preclassic), the lowest date from these fields, through about 1290 BP, during the Late Classic, when the Maya were digging a large number of ditches in this environment and piling up

Fig. 10. Graphs of magnetic susceptibility (A) and ICP-MS elemental analysis (B) at the coastal plain wetland soil profiles (66 E and C) and magnetic susceptibility (C) at the floodplain soil profile, Cacao Creek, Belize.

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the ditched sediments. Within the gypsum layers in most profiles is a faint paleosol at 0.5–0.8 m, a zone that has two radiocarbon assays: 1300–1490 BP and 440–50 BP and pollen indicative of agriculture (Zea mays and Persea americana) (Beach et al., 2008). The modern topsoils are generally cumulic, highly melanized, and have an elevated concentration of metals on a par with the buried paleosols, indicative of long stable formation after the abandonment of these fields, likely in the Maya Terminal Classic (c. 1100 BP). The Rio Bravo provides an example of a perennial stream with widely fluctuating flows occupying a structural valley within the karst uplands of Belize (Fig. 1). In the only work on these systems, Beach et al. (2003) described four soil pits across the floodplain in the vicinity of the site of Dos Hombres. Two of the sites near the Rio Bravo had thin topsoils and deep, calcareous clay alluvium with occasional sand lenses to a depth greater than 2 m. Two other soil profiles at greater distance from the Rio Bravo displayed paleosols buried by calcareous sediments and well formed topsoils at 0.85 to 0.9 and at 1.9 m depths, which were datable by ceramics and one radiocarbon date (2155–1710 BP) from charcoal in the top of the paleosol near unidentifiable ceramics. Hence, this thick paleosol was the upper horizon of a Vertisol in the Late Preclassic that was aggraded after this point, though slope instability has declined with the development of the present, thick and clay-rich topsoil (Beach et al., 2003). Ten km downstream in the Rio Bravo floodplain lies the next set of excavations at Cacao Creek, eight km to the south and downstream from the coastal plain sites at Blue Creek (Fig. 1). We excavated eight units in this area to study soil profiles and evidence for the chronology of aggradation in this fluvial environment (Table 5). These soils have clay textures (68–89% fine and coarse clay), topped by thick Oa horizons of 0.3–0.5 m, and nearly as gypsum-rich as the coastal plain depression soils, with gypsum ranging from 45–65%. Profiles here (Fig. 8F) have occasional flood deposits of calcareous sands and paleosols, buried at 0.9–1.15 m and 1.55 m, which lack most the distinct melanization that marks paleosols at Coastal Plain sites. Three new radiocarbon dates from these aggrading wetlands have helped us define rates of landscape aggradation. A date on charred organic matter at 0.5 m is 760–940 BP, another on a paleosol at 0.9–1.1 m is 960–1170 BP, while the deepest paleosol (at 1.55 m) dated to 1310– 1510 BP. These faintly melanized paleosols were short-term surfaces, but several lines of evidence point to them as paleosols, including increases in magnetic susceptibility, organic matter, artifacts, and phosphorous (Fig. 8F; Table 5). Evidence from a carbon isotopic ratio (δ13C) profile for Cacao Creek also suggests large scale change in these wetlands and provides a comparison of this well dated profile with the other carbon isotopic ratios discussed above. At Cacao Creek, δ13C ranges from little input of C4 species at 0.45 m under recent tropical forest and herbaceous wetland vegetation to significantly increased inputs from C4 vegetation into soil organic matter in the Maya Classic layers. The δ13C increases by more than 7 ppt through this profile from −29.61 at 0.45 m to −22.5 in Classic era sediments at 1.2 m (Fig. 4D). Not surprisingly, sedimentation has been faster in these floodplains than the coastal plain sites. The depression wetlands had average rates of aggradation of about 0.5 mm year− 1 compared with 1 mm year− 1 on this active floodplain, based on radiocarbon dates of the profiles. For comparison, another study of wetland formation in northern Belize estimated Histosol formation of about 1 mm year−1 (Kim and Rejmánková, 2002). The floodplain sediments, compared with those of the depressions, are more carbonate rich than gypsum rich, are finer textured, more alkaline, and have 2 to 4 times less magnesium. These differences reflect that the floodplain sites receive more varied water sources from a much wider area. Micromorphology of these soils reveals the floodplain sediments aggraded from more of an even mix of fluvial clays and silts deposited during overbank flow and groundwater precipitation of gypsum. In contrast, sediments in nearby karst depressions are more uniformly gypsic.


7.7. Northern coastal plain, Belize The initial research that recognized regional paleosols as evidence for aggradation in relation to Holocene sea level rise focused on Maya archaeology and the questions of ancient Maya agriculture (Turner and Harrison, 1983; Pohl, 1990; Furley et al., 1995; Pohl et al., 1996). These research teams studied wetland agriculture in northern Belize, and found buried soils under the rectilinear surface expression of wetlands. The sites studied in the northern coastal plain of Belize lie 40–50 km north of the Blue Creek wetlands along the Rio Hondo and the New River (Fig. 1) and tributaries and at about 2 masl. These sites at Douglas, Cobweb, and Pulltrouser swamps all had paleosol surfaces composed of mineral and organic soils that dated to the Archaic and Preclassic, and have since been buried by 1–2 m of sediment (Pohl et al., 1996) (Table 6). In most cases ancient Maya canals extended through the paleosols, and Pohl et al. (1996) argued that Preclassic ancient Maya farmers used this soil surface during a regression or still stand of sea level rise around 3500 BP. These paleosols were buried by alluvial carbonate rich and evaporitic gypsum rich deposits of Maya Clay from slope erosion after this time period through about 1100 BP (Pohl et al., 1996; Pope et al., 1996). Another wetland field site, Sierra de Agua, 40 km south of Blue Creek along a tributary of the New River produced a similar record of aggradation (Beach et al., 2003). Here an equally well developed paleosol was buried by fluvial deposition and gypsum precipitation, and charcoal from the Ab horizon parallels the dating of the Chan Cahal buried soils (3440–2890 BP, Table 6). Moreover, most of these coastal plain wetland sites have comparable soil profiles to those at Blue Creek. Most dated paleosols have abundant evidence of Preclassic Maya use and are broadly of the same age (formed by 3000 BP), have since been buried by up to 2 m of alluvial and evaporitic sediments around 3000–1000 BP, and most recently been capped by thick organic and mineral topsoils. The soils in the northern coastal plain, however, started to aggrade earlier, which may be related to lower elevation and sensitivity to the impact of sea level rise. 8. Mechanisms of aggradation Numerous studies exist on the rates and processes of aggrading watersheds and most of the evidence comes from sediment deposition from a dated sequence and often above a baseline paleosol (Beach, 1994; Knox, 2006). A paleosol represents a relatively steady surface for pedogenosis to occur and the overlying sediments indicate change from this equilibrium that is too fast for cumulic soil formation or another episode of pedogenesis. The proximate mechanisms that could cause this are an increase in watershed erosion and deposition or an increase simply in deposition over conveyance of sediments elsewhere. The ultimate causes for increased erosion over time could be changes in parent material (through erosion or weathering), vegetation changes driven by human or natural agency, climate changes, base level changes, stream capture, and high magnitude and low frequency events. The main causes for increased deposition in a basin are elevated base level (rising sea levels), rising water tables, change of a water source to one with high dissolved loads, human manipulation of sinks, and stream capture. Volcanic ejecta and mass wasting are certainly important in the Highlands (for example in El Salvador, Dull, 2007), but little scientific literature suggests significant quantities from these agents, although these, as well as aeolian deposition, contribute small amounts to regional sediments (Gunn et al., 2002a; Muhs et al., 2007). Differentiating human induced from climatic and other natural processes is always challenging. For example, at Lake Patzcuaro, Mexico in the Mesoamerican Highlands, two studies came to different conclusions about the timing and drivers of erosion. Based on estimates of lake sedimentation, O'Hara et al. (1993) argued that human-induced soil erosion was extremely high in three


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Precolumbian surges, the last with high populations and in spite of the ancient soil conservation. Fisher et al. (2003), however, argued from soil cores, exposures, and trenches that most erosion occurred in two stages: first during the initial stages of disturbance for site construction (1200–1800 BP) and second in the Colonial period with population desertion. For ancient Maya lowlands landscapes, the erosion record thus far has two main phases of accelerated erosion: a Preclassic pulse (3500– 1700 BP) and Classic pulse (1700–1100 BP). Multiple records of a Preclassic pulse occur in a wide swath from El Salvador to the bajos and lakes of the central Petén and to far northern Belize. Two lines of reasoning buttress natural processes as factors for increased erosion: the Archaic and Preclassic had the highest climatic instability based on the Cariaco cores (Haug et al., 2001) and the Preclassic populations were small and had only meager, stone tools. But three other lines of evidence support human causes. First, populations and land uses were increasing in the Preclassic. Second, agriculture was spreading during this time into new lands and we have sparse evidence for conservation at this time. Third, rates of erosion in the study by Anselmetti et al. (2007) were highest in the Late Preclassic but remained twenty times higher than pre-Maya rates during the Classic period, coinciding with the highest intensification of land use. The decline in rates of sedimentation in the Classic period runs counter to a Malthusian explanation for erosion that ties population directly to degradation, but not an explanation that recognizes that more intensive landuse is not mutually exclusive of soil conservation (Beach et al., 2006b; Dunning et al., in press). Indeed, the Classic period decline of rates of sedimentation comes with increased conservation over most the region and source reduction as the thin Rendoll soils became more truncated and gullied, which may have limited sediment-supply for sedimentation. Moreover, sediments deposited in the El Salvador and Petén lakes and bajos from this period comes with greater quantities of charcoal, economic and disturbance pollen, influxes of phosphorus and some artifacts. The evidence is, thus, stronger for human induced erosion in many environments, especially at sites such as Cancuén and Copán where erosion increased later and coincided with the most intensive land use and little evidence for local soil conservation (Beach et al., 2003, 2006a). At several sites, such as Cancuén and the Petexbatún, ancient erosion and aggradation were significant even with gentle slopes. Sedimentation declined sharply in most sites with the diminution of the Terminal Classic, but then increased with intensified land use in recent times (Beach et al., 2006b).

Evidence for large floods comes from two bajos and one large area of coastal plain wetlands, and all date between 2300 and 1600 BP. These two events made up about half the Late Holocene sediment in the two bajos (Fig. 5C), and these dates overlap the coastal plain event that makes up about 7% of Late Holocene sediment here. Recent hurricanes have left severe impacts, and we need more studies of the long-term recurrence intervals. Many basins in the coastal plain and bajos also had significant human soil manipulation for agriculture, transportation, and defense. Most of this anthropogenic manipulation did not contribute to aggradation but occurred in previously aggraded sediments. The evidence we have discussed thus far explains the aggraded sediments that are carbonate and silicate clay rich and, thus, similar to sediments of the upland slopes. But sediments from the Belize (and adjacent Mexico) northern coastal plain are distinctly gypsic. A wide range of environments in the Rio Hondo, the New River, and other northern river systems, separated by at least 50 km, aggraded by c. 1– 2 m with gypsic sediments above paleosols sometime between 3000 and 1300 BP. Evidence thus far shows several factors contributed to this aggradation, including ancient Maya manipulation of the landscape, soil erosion, and gypsum evaporate formation. Through extensive water chemical analysis, Luzzadder-Beach and Beach (in press) linked the Blue Creek depressions with gypsum aggradation from rising water tables and suggests a model similar to that Pohl et al. (1996) hypothesized for the northern coastal plain sites. This represents a rare geomorphic agent of aggradation, which also occurred on a Pleistocene land surface near Lake Eyre, South Australia (Twidale, 1984). 9. Results of water chemistry This section presents 10 new physical and mineral analyses of 18 regional water samples from 2006 and 2007 to compare sites across the northern coastal plain and test the model of gypsum precipitation across this broader landscape. These revealed significant regional differences between upland and lowland sites, adding to and confirming fifteen years of prior water quality studies in the region (Luzzadder-Beach and Beach, in press) (Table 8). We present these findings south to north from the upland karst basins down through the coastal plain. Electrical Conductivity (EC) ranged from a low of 189.5 μS in the uplands (Chan Chich Creek at Suspension Bridge), to a high of 3310 μS in the lowlands (Four Mile Lagoon), with an average for the uplands of 514 μS and the lowlands of

Table 8 Selected water quality results 2006–2007, northern Belize stations Sample location and site Upland sites 1. R. Azul/Blue Cr. at Rapids 2. Upper Little Chan Chich Cr. 3. Chan Chich Cr. at Susp.Br. 4. R. Bravo at Cedar Crossing 5. Aguada at Rosita Upland means Lowland sites 6. New River at Lamani 7. Laguna Verde, 7 m 8. Birds of Paradise I Canal 9. R. Bravo below Cacao Cr. 10. Cacao Cr. at BP Trailhead 11. Lagoon SW of BP 12. Rio Hondo at S. Antonio 13. Cenote at Albion Isl. Quarry 14. Four Mile Lagoon nr. Chetumal 15. Hector Cr. nr. Belize City Lowland means


NO−3-N− mg L− 1

Ca2+ mg L− 1

Mg2+ mg L− 1

Cl− mg L− 1

−1 SO2− 4 mg L

634 496 579 647 216 514

0.4 0.3 0.8 1.3 1.0 0.8

76.8 65.6 59.2 60.8 40.0 60.5

32.6 21.6 39.4 51.4 3.8 29.8

170.7 113.8 170.7 170.7 56.9 136.6

27 6 7 6 6 10.4

682 2560 2050 2130 1071 1043 1851 1010 3310 2840 1855

0.2 0.0 1.2 0.0 0.5 0.0 0.0 1.4 0.0 0.3 0.3

132.8 528 492.8 655.2 201.6 176 336 201.6 387.2 201.6 331.3

13.4 288 208.3 30.2 377.3 196.8 221.7 262.1 282.2 201.6 208.2

170.7 739.7 569 625.9 284.5 284.5 512.1 284.5 1024.2 853.5 534.9

260 1239 1027 798 252 294 810 315 819 147 596.1

TDS mg L− 1

Sal 0/00

Na+ mg L− 1

Alkalinity mg L− 1

304 235 277 309 101 245

0.3 0.2 0.3 0.3 0.1 0.2

124.2 70.7 114.9 112 38 92

328 240 312 360 112 270

328 254 312 366 116 275

325 1240 1810 1070 518 508 896 484 1690 1430 997.1

0.3 1.3 1.0 1.1 0.5 0.5 0.9 0.5 1.8 1.5 0.94

127.2 0 7.8 132.8 0 0 14 0 155 52.1 48.9

156 166 248 348 388 380 228 156 180 92 234

388 2520 2100 1764 2076 1260 1764 1596 2144 1344 118

Total hardness mg L− 1 as CaCO3

T. Beach et al. / Geomorphology 101 (2008) 308–331

1855 μS. This parameter is a measure of total minerals, and is closely related to TDS (for more details see Luzzadder-Beach 1997 and 2000). TDS ranged from 101 to 309 mg L− 1 in the uplands (mean = 245 mg L− 1), and 325 to 1810 mg L− 1 in the lowlands (mean = 997.1 mg L− 1). Salinity ranged from 0.1 to 0.3 ppt in the uplands (mean = 0.2 ppt), and from 0.3 to 1.8 in the lowlands (mean = 0.94 ppt). Although we studied nitrate (NO−3 as N) for another paper on water quality (Luzzadder-Beach and Beach, in press), it was not relevant to aggradation. Sulfate (SO−4 2) ranged from very low values in the upland water sources (6 to 27 mg L− 1; mean of 10.4 mg L− 1) to very high values in the lowland sites (147 to 1239 mg L− 1, mean value of 596.1 mg L− 1) (Table 8). This reflects very different characters of the upland versus lowland waters and the water sources. When combined with calcium (Ca2+), sulfate becomes gypsum (CaSO4·2H2O), which we have observed precipitating in some environments (Luzzadder-Beach and Beach, in press). The threshold saturation value of SO−4 2 is 1600 mg L− 1, (Hounslow, 1995, p. 53), and only one sample from 2007 approached saturation: Laguna Verde at 1239 mg L− 1 in this Wet Season field work (Table 8). Calcium and magnesium are components of total hardness, and the proportional presence varies with hydrogeologic sources. Like sulfate, concentrations of calcium (Ca2+) and magnesium (Mg2+) were five to seven times higher on average in lowland waters compared with upland waters (Table 8). Calcium in upland waters ranged from 40 to 76.8 mg L− 1, with a mean concentration of 60.5 mg L− 1. Calcium levels of the lowland waters ran from 132.8 to 655.2 mg L− 1, with a mean concentration of 331.2 mg L− 1. One site, the Rio Bravo below Cacao Creek (655.2 mg L− 1), was above the calcium saturation threshold of 636 mg L− 1 (Hounslow, 1995, p. 53; Luzzadder-Beach and Beach, in press). Magnesium ranged from 3.8 to 51.4 mg L− 1 in the uplands (mean = 29.8 mg L− 1), to between 13.4 and 377.3 mg L− 1 in the lowlands (mean = 208.2 mg L− 1). These are very common ions in carbonate or evaporite systems, often deriving from dolomite and limestone dissolution (Langmuir, 1997), and together they form the dominant ions of total hardness in these waters. Mean total hardness (measured as CaCO3, mg L− 1) was 275 mg L− 1 in the uplands, and 1186 mg L− 1 in the lowlands. Chloride (Cl−) levels were fairly high in the lowland sites (ranging from 170.7 to 1024.2 mg L− 1, mean = 534.9 mg L− 1; nine samples exceeded the 250 mg L− 1 standard) and were fairly low in the upland sites (ranging from 56.9 to 170.7 mg L− 1, mean = 136.6 mg L− 1). Overall, sodium (Na+) levels were fairly low in this region. Upland waters had a mean value of 92 mg L− 1, and lowland waters had a mean value of 48.9 mg L− 1, and are, therefore, not a factor for sedimentation. 9.1. Water chemistry: discussion The presence of sulfate and calcium as dominant ions, especially in the lowland waters, indicates an aquifer source rich in gypsum (CaSO4·2H2O). This corroborates fifteen years of prior findings of the lowland waters in the region being dominated by, and in some cases nearly saturated with, these ions in many coastal plain locales (Luzzadder-Beach and Beach, in press). Based on comparing the relative amounts of SO−4 2 to Ca+2and Mg2+, we can also interpret different sources for these minerals in the upland and lowland waters. In the upland waters, distinctly more Ca+ 2 (mean = 60.5 mg L− 1) and Mg+ 2 (mean = 29.8 mg L− 1) occurs than SO−4 2 (mean = 10.4 mg L− 1); therefore, the likely geologic sources are from a calcite or dolomite, and not likely a gypsum source (Hounslow, 1995). In the lowlands, less Ca+ 2 (mean 331.3 mg L− 1) and Mg+ 2 (mean = 208.2 mg L− 1) occurs than SO−4 2 (mean = 534.9 mg L− 1), meaning that, as the amounts become more equal a gypsum source exists, or as the sulfate becomes more dominant the calcium is being removed, likely by calcite precipitation in this environment (Hounslow,1995). Because this entire region has medium alkalinities (upland mean = 270 mg L− 1, the lowland mean is 234 mg L− 1), it is less likely that the process of pyrite oxidation or


natural softening are removing calcium (Hounslow, 1995). In some cases, sulfate and calcium are precipitating out from saturation as surface water evaporates, observed in the field and in samples by microscopy (Luzzadder-Beach and Beach, in press). In addition to the high concentrations of minerals discussed above, the proportions of ions (in percent, mass balance of cations and anions equivalents per million (epm)) were also analyzed to characterize the waters, their sources, and possible mixing of these waters. The upland and lowland waters grouped separately. Lowland waters were generally strongly calcium, magnesium, and sulfate in nature, and are considered in the area of permanent hardness (Fetter, 1994, p. 423, and Hounslow, 1995, p. 89). Upland waters were characterized as more balanced in nature, and on the side of temporary hardness (Fetter, 1994, p. 423; Hounslow, 1995, p. 89) as calcium, magnesium, and carbonate/bicarbonate (medium alkalinities) with lower chlorides (mean = 136.6 mg L− 1). Minimal mixing of these two groups occurs, but within each respective group mixing appears to happen. One of the more interesting findings provides an exception to the lowland water chemistry: despite its high ion concentrations, the New River at Lamanai exhibits an overall character (proportional distribution of ions) that is closest to all of the upland sources where its headwaters reside, rather than to the other lowland water sources, indicating little mixing of lowland waters with the New River system at this location. Hector Creek, in the lower Belize River Valley of central Belize, was included as a spot check in this study, and its character was much more like the other lowland sources (on the coastal plain) rather than the upland sources. Three water sources of the 18 samples were outliers: the cenote at Albion Island, the Lagoon near Cacao Creek, and another arm of Cacao Creek. All three sources showed some imbalance or error in the mass balance of cations and anions. Even after we repeated these measurements, they displayed higher proportions of Ca+ 2 and Mg+ 2 than expected based on the comparison with the known, opposing anion mass (alkalinity, sulfate, and chloride as epm). Ultimately these waters characterized closely with the other lowland water sources. The imbalance may result from evaporation or super saturation of carbonate minerals (Langmuir, 1997) or from the presence of an as yet unknown anion beyond the scope of this study. In short, water chemistry across this region clearly indicates the large and abrupt geographic divide between upland waters, with low ionic contents, and lowland waters, with high sulfate and calcium concentrations. 10. Conclusions Evidence of geomorphic change across the central Maya Lowlands as well as the broader Maya realm is patchy. We have a growing opus of karst studies and studies of soils and paleoecology of the karst sinks and wetlands of the region, but we still know little about the geomorphic history of most river systems. We also know particularly little about several important sub-regions within this broader region, such as the Maya Mountains of Belize, the Rio Motagua of Guatemala, and the Usumacinta watershed of Petén, Guatemala and Chiapas, Mexico. The available evidence for slope erosion is still limited to a few observations and localized measurements, often from sites that have experienced long histories of human disturbance. Several local reports exist for Belize and the Petén of Guatemala of high soil losses in recent decades, which provide a modern analog to compare with ancient erosion on these moderately sloping landscapes. The landscape today may still preserve its history of erosion in situ: many soil profiles from the slopes of the region display thin Rendoll soils. But we have too few studies of the age of soils along a variety of slopes or buried erosion features that might adequately identify the timing and processes of erosion. Thus, we still cannot fully judge hypotheses that the thin, young soils of the Petén are legacies of ancient Maya erosion. The main evidence for erosion during the Maya period is aggradation above some baseline: paleosols, stratigraphy, artifacts,


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or structures. The other lines of evidence are the clues that come washed in with the sediments: artifacts, ecofacts, dateable materials, and sediment chemistry. Because sediments can deposit from the suspended or bedload or precipitate from dissolved load, the chemistry of source water is another important line of evidence that is often overlooked. The best studied environments thus far are the Petén Lakes and adjacent bajo sinks. The original model of aggradation of ‘Maya Clays’ during the Maya Period from 3000 to 1000 BP still holds in most of the region, but we are refining this sequence in lakes and bajos. In the Petén lakes soil erosion accelerated early (by 3000 BP) and reached a peak around 2000 BP in the Maya Late Preclassic, after which it declined by half (still 20 times higher than background rates) during the Maya Classic up to 1100 BP before dropping back to background levels until recent times. The history of erosion from bajos is more complicated, because of post-depositional disturbance and to variations in proximal–distal deposition. Although several bajos have buried paleosols and paleochannels that date from the Pleistocene, many more have Archaic to Preclassic paleosols buried by 1–2 m. In addition, those sequences we have investigated often display evidence of pedogenesis between these Preclassic paleosols and the modern topsoil. Evidence here suggests earlier aggradation (Preclassic) as well as the expected Classic aggradation. Some karst sinks, such as at Cancuén, like some river valleys, have mainly experienced aggradation in the Late Classic; others like the bajos around San Bartolo have mainly Preclassic aggradation; and yet still other sites like the Dumbell Bajo show little evidence of aggradation. Most of these sinks remained stable or depositional until the later mid Holocene. The alluvial fans sequences in our excavations, for example, were largely stable environments in the mid Holocene, and then begun to aggrade through to around 1000 BP, with either stable or slow sediment accumulation in the last millennium. One alluvial fan and a nearby karst sink at Blue Creek exemplify this pattern: repeated instability in the Archaic through Classic Period but stable pedogenesis over the last 1000 years. Research to date indicates that river valleys were largely stable until the Late Holocene. River systems would have had to adjust to higher base levels and transgressing seas into the Middle Holocene, after which sea level rise became more stable. Aggradation in the river valleys started at about 3000 BP (Maya Preclassic) in many floodplain sites and continued through the Classic period, though research has focused on the Classic period mainly in several areas, such as the Motagua River area near the Maya site of Quiriguá. In the Rio Bravo, on the other hand, Preclassic paleosols are buried by Preclassic and Classic period sediments with only sporadic sand lenses and covered with thick topsoils, showing recent pedogenesis and relative stability. The Copán River was a rare example of incision creating a new floodplain and terrace after the major Maya abandonment from 1000–500 BP. If, after large scale abandonment and reforestation, suspended sediment loads decreased with reforestation, then we may expect to find more evidence of incision in river valleys from this region, but the continued slow rise of sea level and lowered runoff could have compensated for this. Several sites, Copán, Quiriguá, and the Xibun, experienced higher energy, sand and gravel, depositional events but the dating has been inexact. Too little research exists to compare magnitude and frequency of geomorphic events and processes in the Holocene. The wetland depressions and floodplain sites of the coastal plain region have, like the broader regional floodplains and bajos, 1–2 m of aggradation, but with distinct Preclassic and Late Classic phases, and quite different chemical natures. The first phase occurred as wetland and floodplain sites in the far northern coastal Plain of Belize started to aggrade soon after 3500 BP with sea level rise induced inundation and organic soils formation. The second phase started later, from the Preclassic through the Late Classic as ‘Maya Clays’ and gypsic evaporites buried the landscape. This model supports our findings from the Blue Creek wetlands on the southern end of Belize's northern

coastal plain. Here Archaic and Preclassic paleosols also became buried by Preclassic through Late Classic organic matter, clay, and evaporite accumulation. We have identified five main mechanisms that may be responsible for the aggradation events of the late Holocene: accelerated erosion by extreme climatic instability and human land use change, extreme floods, human manipulation of depression soils, and sea level rise induced uplift of nearly saturated water tables. Another possible contributor to erosion and aggradation was accelerated runoff from widespread deforestation, which may have hastened all of the other factors. The evidence for significant climatic instability precedes the start of and declines before the end of Preclassic aggradation. Periods of climatic instability also overlap the millennia of human induced landuse changes, but the aggradation evidence more precisely parallels intensive land uses than the climate evidence. For example, aggradation continues for 1000 years beyond climatic instability, tracking the Preclassic through to the Classic Maya periods, though aggradation declines by half through the Late Classic period when land use intensities would have been highest. Several lines of evidence may explain this decreased aggradation: the decrease is still much higher than background rates, reduction of sediment supply from upland slopes truncated by more than a millennium of land uses, periods of reduced urban growth, less climatic instability, and the widespread growth of conservation, preserved partially in stone terraces dating mostly to the Classic period. A third factor in two bajos and a region of the coastal plain were high energy events that occurred between 2200 and 1600 BP in the bajos and between 2300 and 2000 BP in the coastal plain wetlands. The high magnitude events made up about half the Holocene sediment in these two bajos but only about 7% in the coastal plain area. Ample evidence also exists for ancient Maya burying landscapes consciously to reclaim wetlands and build terraces and channels, but this led to only a small magnitude of aggradation. The majority of evidence, reconstructed from sediment chemistry, artifacts, ecofacts, and from modern analogs, shows that human induced deforestation and erosion was the main cause of accelerated erosion and aggradation in all sinks except those located on the coastal plain floodplains and depressions. Here the aggraded sediments are dominantly gypsum rich sediments with thin Histosols and short-term paleosols, rather than the carbonate and silicate clays that we found in the upland sediments and slope soils. This fourth mechanism of aggradation corresponds to sea level rise inducing wetland and peat formation and forcing up water tables nearly saturated with sulfate and calcium. The sediment chemistry parallels the groundwater chemistry over a 100 km expanse across the coastal plain to the sea. Eroded sediment and human manipulation contributed to this aggradation but because gypsum, which comes from ground water, overshadows the mineralogy, evaporite formation from rising water tables is the probable explanation for coastal plain aggradation. Thus, aggradation has several causes in this region, but the two main drivers are different ones for adjacent geomorphic landscapes, though they are deceptively similar to the naked eye. Acknowledgements We thank the following organizations for funding and support: Georgetown University's School of Foreign Service and Graduate School; The National Geographic Society's Committee for Research and Exploration (Grant Nos. CRE-7506-03, CRE-7861-0s5; T. Beach and S. Luzzadder-Beach PIs), The National Science Foundation (Grant No. SBR-963-1024, V. Scarborough and N. Dunning PIs; Grant Nos. BCS-0241757 and BCS-0650393, N. Dunning PI); George Mason University's Center for Global Studies and Provost's Office; The Smithsonian Center for Materials Research and Education (SCMRE); and The University of Cincinnati. We recovered data from Belize as part of the Maya Research Program, directed by Jon Lohse and Tom

T. Beach et al. / Geomorphology 101 (2008) 308–331

Guderjan, and the Programme for Belize Archaeological Project, directed by Fred Valdez Jr., and with the gracious cooperation of the Department of Archaeology, Ministry of Tourism and the Environment, and the Programme for Belize. Data from the Bajo Zocotzal near Tikal and the Bajo La Justa near Yaxha derives from the Subproyecto Intersitios, directed by Vilma Fialko, as part of the larger Proyecto Triangulo of the Instituto de Antropología e Historia (Guatemala), directed by Oscar Quintana. Funding for Dunning's work was provided by a grant from NASA to Tom Sever, Pat Culbert, and Vilma Fialko. Rosa Maria Chan directed excavations at the Bajo Zocotzal. Dunning's work near San Bartolo, Guatemala was part of the Proyecto San Bartolo, directed by William Saturno and Mònica Urquizu. These investigations also included three University of Cincinnati geoarachaeology field courses, and their participants are thanked for producing some of the data reported here. David Klemm, Georgetown University Medical School, helped develop and draft many of our graphics. We wish to thank the following individuals for their contributions to these investigations: Drs. John Jones, Steve Bozarth, Tom Garrison, Rob Griffin, Paul Hughbanks, Laura Kozikowski, Julie Kunen, Kevin Pope, Vernon Scarborough, Kerry Sagebiel, Steve Houston, Richard Terry, and Kim Cox, Esq.; and fifty Georgetown, George Mason, and Cincinnati graduate and undergraduate students. The communities of Blue Creek and San Felipe have been gracious hosts and field work partners for our research teams over the years and deserve many thanks. Michael Day and several anonymous reviewers have our gratitude for their excellent suggestions to improve this manuscript. The views expressed herein and any errors or omissions are ours alone. References Abrams, E.M., Rue, D.J., 1988. The causes and consequences of deforestation among the prehistoric Maya. Human Ecology 16, 377–395. Anselmetti, F.S., Ariztegui, D., Brenner, M., Hodell, D., Rosenmeier, M.F., 2007. Quantification of soil erosion rates related to ancient Maya deforestation. Geology 35, 915–918. Antonie, P., Skarie, R.L., Bloom, P.R., 1982. The origin of raised fields near San Antonio, Belize: an alternative hypothesis. In: Flannery, K. (Ed.), Maya subsistence: Studies in memory of Dennis E. Puleston. Academic Press, NY, pp. 227–238. Ashmore, W., 2007. Settlement Archaeology at Quiriguá, Guatemala. University of Pennsylvania Museum of Archaeology and Anthropology, Philadelphia. Beach, T., 1994. The fate of eroded soil: sediment sinks and sediment budgets of agrarian landscapes in southern Minnesota, 1851–1988. Annals of the Association of American Geographers 84, 5–28. Beach, T., 1998a. Soil catenas, tropical deforestation, and ancient and contemporary soil erosion in the Petén, Guatemala. Physical Geography 19, 378–405. Beach, T., 1998b. Soil constraints on Northwest Yucatán, Mexico: pedoarchaeology and Maya subsistence at Chunchucmil. Geoarchaeology 13 (8), 759–791. Beach, T., Dunning, N.P., 1995. Ancient Maya terracing and modern conservation in the Petén Rain Forest of Guatemala. Journal of Soil and Water Conservation 50, 138–145. Beach, T., Luzzadder-Beach, S., (in press). Geoarchaeology and plain aggradation near Kinet Höyük, Eastern Mediterranean, Turkey. Geomorphology. doi:10.1016/j. geomorph.2008.03.006. Beach, T., Luzzadder-Beach, S., Dunning, N., Hageman, J., Lohse, J., 2002. Upland agriculture in the Maya Lowlands: ancient Maya soil conservation in northwestern Belize. The Geographical Review 92, 372–397. Beach, T., Luzzadder-Beach, S., Dunning, N., Scarborough, V., 2003. Depression soils in the lowland tropics of northwestern Belize. In: Gómez-Pompa, A., Allen, M., Fedick, S.L., Jiménez-Osornio, J.J. (Eds.), Lowland Maya area: three millennia at the humanwildland interface. 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