Connecticut, Maine, Massachusetts, New Hampshire, New York, Rhode Island, Vermont
HA 730-M

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Aquifers are present in consolidated rocks in the Central Lowland, the St. Lawrence Valley, the Adirondack, and the New England Physiographic Provinces in Segment 12. These aquifers are in three principal types of rocks: carbonate rocks

(fig. 85), sandstone, and crystalline rocks of igneous or metamorphic origin. No regional study has been conducted for aquifers in any of the three principal rock types in the segment; accordingly, examples of smaller scale studies are described in this report. The examples were chosen because they represent ground-water occurrence, movement, and quality for each principal rock type.

Consolidated-rock, or bedrock, aquifers also are present in the Appalachian Plateaus, the Valley and Ridge, and the Piedmont Physiographic Provinces in Segment 12. The bedrock aquifers in these provinces are carbonate-rock and sandstone aquifers that generally yield only small volumes of water and are of local extent. The bedrock aquifers in the Appalachian Plateaus, the Valley and Ridge, and Piedmont Provinces are not described in this report.


The aquifers in the Lake Erie-Niagara River Basin are an example of carbonate-rock aquifers that are characteristic of the Central Lowland Province of western New York. The aquifers are in the northern one-third of the basin and consist of flat-lying limestone and dolomite with some gypsum and abundant interbedded shale. The aquifers typically yield only small to moderate quantities of water to wells; the water is hard, and saltwater is present in places, commonly at shallow depths.

The Lake Erie-Niagara River Basin in western New York is bordered by Lake Erie and the Niagara River on the west and extends eastward to about the middle of Genesee County (fig. 86). The basin includes the area near the city of Niagara Falls in which streams drain to the Niagara River. A strip of land bordering Lake Erie is part of the Eastern Lake Section of the Central Lowland Physiographic Province, which is a region of low relief. The remainder of the basin is in the rugged Appalachian Plateaus Physiographic Province, which has considerable relief, but all the aquifers described are in the Central Lowland Province.

The basin has been glaciated, a process that scoured the bedrock surface and left a veneer of till in the upland areas and thick, complex valley-fill deposits that consist of ice-contact, outwash, and glacial-lake deposits in the deeply eroded bedrock valleys. The valley-fill deposits form principal aquifers that are hydraulically connected with bedrock aquifers and with streams that traverse the surface of the valley fill.

Bedrock in the basin consists chiefly of limestone, dolomite, and shale. The carbonate rocks and the shale are virtually impermeable as homogeneous rock. These rocks, however, have been subjected to regional tectonic stresses and are vertically and horizontally fractured. The fractures provide openings for the storage and transmission of water. Fracture permeability is enhanced in limestone and, to a lesser extent, in dolomite by dissolution of the rock by ground water. A similar enhancement of permeability is produced by dissolution of interbedded gypsum in some rock units.

The principal bedrock aquifers in the Lake Erie-Niagara River Basin are (1) a limestone aquifer that consists of the Onondaga Limestone, the Akron Dolomite, and the Bertie Limestone; (2) the Camillus aquifer, which consists of the Camillus Shale, the Syracuse Formation, and the Vernon Shale; and (3) the Lockport aquifer, which consists of the Lockport Dolomite. These aquifers differ considerably in water-yielding characteristics but generally yield only small to moderate quantities of water to wells. The aquifers, however, are typical of bedrock aquifers in the Central Lowland Province of New York and are presented as an example of aquifers in that hydrologic environment.


Bedrock in the Lake Erie-Niagara River Basin consists chiefly of stratified limestone, dolomite, and shale of marine origin. The distribution of these units at the bedrock surface is shown in figure 86, and their lithology and vertical sequence are shown and described in figure 87. The units are Silurian and Devonian in age and are virtually flat lying, with a gentle dip to the south of only about 30 to 40 feet per mile (fig. 88).

The bedrock surface was deeply eroded by weathering and stream action prior to glaciation and by glacial scour during glaciation. The surface increases in altitude to the south, in contrast to the dip of the rocks. Thus, rocks that form the bedrock surface in parallel, east-northeast-trending belts are successively younger toward the south (fig. 86).

Much of the bedrock in the Lake Erie-Niagara River Basin consists of black to gray carbonaceous shale with minor calcareous beds and limestone layers (fig. 87). The exceptions are the five formations, which form the three aquifers, that consist of limestone, dolomite, and gypsiferous shale primarily of Silurian age.

The Lockport Dolomite is the lowermost carbonate-rock unit and overlies the Rochester Shale. The Lockport forms the bedrock surface in the northern part of the basin (fig. 86) and consists mainly of fine- to coarse-grained dolomite. Gypsum is present as nodules along some bedding-plane surfaces in the Lockport. The maximum thickness of the Lockport is about 150 feet. Near the base of the Lockport, the formation is divided into the Decew Dolomite Member and the overlying Gasport Limestone Member.

The Vernon Shale, the Syracuse Formation, and the Camillus Shale overlie the Lockport Dolomite and form the bedrock surface to the south of the Lockport (fig. 86). Outcrops of these formations are rare because of the low relief of the area and the cover of glacial deposits. The three formations consist chiefly of shale; however, considerable massive mudstone, limestone, and dolomite are interbedded with the shale (fig. 87). Gypsum is present in beds as much as 5 feet thick, as well as in thin lenses and veins.

A sequence of limestone units-the Bertie Limestone, the Akron Dolomite, and the Onondaga Limestone (fig. 87)-overlies the Camillus Shale. The Bertie Limestone and the Akron Dolomite are of Silurian age and are separated from the overlying Onondaga Limestone of Devonian age by an uncon-formity or erosional contact.

The Bertie Limestone consists mostly of dolomite and dolomitic limestone with interbedded shale particularly in the lower parts (fig. 89). The middle part of the Bertie is a massive dolomite and dolomitic limestone. The upper part is gray dolomite and shale with beds of variable thickness. Its maximum thickness is about 55 feet.

The Akron Dolomite is a fine-grained dolomite with beds that vary in thickness from a few inches to about 1 foot. The upper contact of the Akron is erosional and generally is marked by remnants of shallow stream channels that contain sandy lenses. The thickness of the Akron generally is between 7 and 9 feet.

The Onondaga Limestone, which has a maximum thickness of about 110 feet, forms the upper two-thirds of the carbonate-rock sequence. The Onondaga consists of three lithologies: a lower coarse-textured, crinoidal limestone that is about 10 feet thick; a middle cherty limestone that is about 45 feet thick; and an upper limestone that is about 55 feet thick.

The carbonate-rock sequence is overlain by the Marcellus Shale (fig. 87), which is a gray to black fissile shale with a basal, 10-foot-thick limestone sequence; the thickness of the formation ranges from 30 to 50 feet. The Marcellus Shale is, in turn, overlain by the Skaneateles Shale.


The bedrock units in the northern part of Lake Erie-Niagara River Basin have been divided into three aquifers. In descending order they are the limestone, the Camillus, and the Lockport aquifers. The aquifers are not separated by confining units, but can be distinguished by their contrasting water-yielding characteristics. Except where they form the bedrock surface and are mostly covered by glacial deposits, the three aquifers are overlain and underlain by thick sequences of shale that form effective confining units.

The limestone aquifer, which consists of the Onondaga Limestone, the Akron Dolomite, and the Bertie Limestone (fig. 89), yields water mostly from solutionally enlarged fractures, bedding planes, and other openings in the rock. Bedding planes or horizontal fractures typically are the most enlarged and important water conduits; however, vertical fractures conduct some water between horizontal openings.

The area where the limestone aquifer forms the bedrock surface is drained by Tonawanda Creek and its major tributaries (fig. 86). Water enters the aquifer in the interstream areas by infiltration into joints and fractures. Some of the water is discharged laterally to the streams, and some percolates downward into the Camillus Shale.

The transmissivity of the limestone aquifer ranges from 535 to 3,340 feet squared per day. Transmissivity and specific capacity data are summarized in table 12. A number of large-yield wells in Buffalo, Cheektowaga, Williamsville, Pembroke, and Batavia are completed in the limestone aquifer. These wells yield as much as 300 gallons per minute from the Akron Dolomite and the Bertie Limestone. Most wells completed in the aquifer, however, commonly yield about 30 gallons per minute.

The Camillus aquifer is by far the most productive aquifer in the basin. Industrial wells completed in the aquifer in the vicinity of Buffalo and Tonawanda yield from 300 to 1,200 gallons per minute, and large volumes of water from the aquifer entering mines elsewhere in the basin indicate that large well yields are possible locally. The Camillus Shale contains extensive interbedded gypsum, which is more soluble than the surrounding shale, limestone, or dolomite. Therefore, as the gypsum dissolves, openings remain that enhance the storage and transport of water. As a result of these openings, the Camillus aquifer has transmissivity values that range from 935 to 9,350 feet squared per day (table 13). The transmissivity, however, is variable vertically and horizontally within the aquifer because of the variability of occurrence of the gypsum and the extent of dissolution.

The Camillus aquifer forms a low topographic trough on its outcrop area, which is traversed by Tonawanda Creek. Water that enters the aquifer discharges mainly to Tonawanda Creek. Other streams that traverse the aquifer, however, are not well incised, and discharge from the Camillus aquifer to these streams is small.

The Lockport aquifer forms the bedrock surface on the north side of the basin. The Lockport Dolomite (fig. 90) forms the Niagara Escarpment and the lip of Niagara Falls. Water-yielding characteristics of the aquifer have been studied in de-tail in the vicinity of the Robert Moses Niagara Powerplant; this area is characteristic of the aquifer.

Horizontal bedding-plane joints or zones of such joints are the principal water-yielding openings in the Lockport aquifer in the Niagara Falls area. Although some water moves through vertical joints and solution cavities from which gypsum has been dissolved, these openings are minor conduits. Water- yielding bedding joints might be present in any stratigraphic horizon; however, those that are areally persistent commonly are in zones of thin beds overlain by thick or massive beds. Seven such areally extensive water-yielding zones have been identified in the Lockport aquifer in the Niagara Falls area (fig. 91). The bedding-plane joints appear to be continuous for miles, but they are not water-yielding everywhere. For example, the water-yielding bedding-plane joint that is about 35 feet above the base of the Lockport aquifer is located in a zone where gypsum nodules along bedding planes have been dissolved, thus providing openings for flow and storage of water. This bedding-plane joint yields some water but is not laterally persistent because of the differential dissolution of gypsum. Outcrop and well log data indicate that similar horizontal conduits are present in other parts of the basin. Where they contain water, they have been widened by dissolution of the dolomite. Some joints stand open as much as one-eighth of an inch. Locally, dissolution along the bedding planes has been sufficient to cause the overlying rock to settle.

An eighth widespread water-yielding zone in the Lockport is a weathered zone in the upper 10 feet of the formation which is extensively fractured (fig. 91). This zone is present only in outcrop areas and is hydraulically connected to the overlying glacial deposits.

A transmissivity of 305 feet squared per day was calculated for the Lockport aquifer on the basis of data collected during dewatering of an 18,000-foot-long powerplant intake conduit near Niagara Falls. This probably is a representative transmissivity value for the aquifer because of the extent of the aquifer involved. Aquifer tests were conducted in the Niagara Falls area by using four wells completed in the aquifer. Transmissivity values of 40 to 135 feet squared per day and coefficients of storage of 0.00001 to 0.0003 were calculated. The smallest values of transmissivity and the smallest coefficients of storage were obtained from wells completed in the lower part of the Lockport aquifer.

The Lockport aquifer is the least productive of the three bedrock aquifers. Yields of wells completed in this aquifer within the Lake Erie-Niagara River Basin range from less than 1 to about 100 gallons per minute. In the Niagara Falls area, yields of wells completed in the lower 40 feet of the Lockport aquifer range from 0.5 to 20 gallons per minute with an average yield of 7 gallons per minute. Wells completed in the upper part of the aquifer yield from 2 to 110 gallons per minute and have an average yield of 31 gallons per minute. Well yields from the Lockport are areally and vertically variable but generally are about 30 gallons per minute.


Ground water is stored in and moves through secondary openings (joints and fractures) in the predominantly carbonate and shale rocks of the Lake Erie-Niagara River Basin. On a regional scale, the rocks can be considered as a single ground-water system with a continuous water table. Water moves through this system under a hydraulic gradient from areas of recharge to areas of discharge.

An idealized diagram that shows ground-water movement in bedrock aquifers of the basin is shown in figure 92. The flow lines describe the theoretical movement of water through the system. The lines of equal hydraulic head, which are at right angles to the flow lines, represent the hydraulic gradient in the aquifer. Hydraulic heads are highest in the upland areas, which are principal areas of recharge, and decrease to a minimum in major stream valleys, which are principal areas of discharge. Thus, in upland areas, hydraulic heads decrease with depth, and flow is downward. In intermediate areas, hydraulic heads are approximately equal in the vertical plane, and flow becomes primarily horizontal. At major stream valleys, hydraulic heads increase with depth, and ground-water flow is upward toward the stream, where it is discharged. The depth of the regional flow system is unknown, but it probably is no more than 300 to 500 feet below land surface because openings in the rocks below those depths are minimal.

Superimposed on the regional flow system are small local flow systems that discharge to secondary (tributary) streams (fig. 92). These systems operate in the same way as the regional flow system, but they are shallow and generally exist entirely within the drainage basin of a tributary stream. Local flow systems are sensitive to droughts during which the water table might decline enough to partly or completely obliterate the flow system and cause the tributary stream to cease flowing.

Inferred ground-water circulation in the Lake Erie-Niagara River Basin is through regional flow systems that have recharge areas in the Appalachian Plateaus Province; the water subsequently moves through the Central Lowland Province and discharges to Tonawanda Creek. The deepest circulating water moves upward toward Tonawanda Creek through joints and fractures in the Camillus Shale and the Lockport Dolomite.

Water movement in the bedrock aquifers has been affected by withdrawals through wells and engineering developments, such as hydroelectric power at Niagara Falls. The movement of water through the upper part of the Lockport aquifer prior to ground-water development is shown in figure 93. Ground water moved from topographically high areas mostly toward the Niagara River downstream from Horseshoe and American Falls. Ground water moved northward along a thin strip adjacent to the Niagara Escarpment. A ground-water divide separated movement toward the escarpment and the Niagara River. Water also moved into the Lockport from the Niagara River upstream from the falls, bypassed the falls, and discharged into the downstream reaches of the river.

The potentiometric surface of the Lockport aquifer changed (fig. 94) after construction of the Lewiston Pump-Storage Reservoir and the Forebay Canal, which conveys water westward from the reservoir through two powerplants to the Niagara River downstream from the falls. An unlined intake conduit also was constructed that trends northward from the Niagara River upstream from the falls to the Forebay Canal. The bottom of the intake conduit is below the water table, and the conduit functions as a line of discharge for the aquifer. The water level in the Lewiston Pump-Storage Reservoir is higher than the water table, and the reservoir functions as an imposed recharge area for the aquifer. These three structures have profoundly rearranged the ground-water flow system of the area.

The intake conduit provides a principal drain for the Lockport aquifer, and water is moving toward the conduit from the east and west for its entire length. Discharge from the reservoir moves into the aquifer and westward toward the conduit. A ground-water divide, which separates the eastward and westward movement, has been established between the conduit and the downstream reaches of the Niagara River because of discharge to the conduit. Water continues to enter the aquifer from the upstream reach of the Niagara River, but some now moves toward the conduit. Some water continues to move through the aquifer and around the falls to the downstream reaches of the Niagara River.


Dissolved constituents in the ground water in the northern part of the Lake Erie-Niagara River Basin are derived primarily from dissolution of the rocks through which the water moves. Water-yielding rocks in the basin contain four soluble minerals: calcite, which is the major constituent of limestone; dolomite; gypsum; and halite, or rock salt. Calcite and dolomite are present throughout the basin especially in the Lockport aquifer and in the limestone aquifer. Most shale formations in the basin, including those of the Camillus aquifer, are calcareous. Calcium, magnesium, and bicarbonate ions are dissolved from the aquifer minerals and contribute to the hardness of the water.

Gypsum is present in the Camillus aquifer and, to a lesser extent, in the Lockport aquifer. The principal dissolutionproducts of gypsum are sulfate and calcium. Halite is present in the Camillus aquifer in the southern one-half of the basin. Dissolution of halite produces sodium and chloride in ground water.

Each of the soluble minerals in the basin also is present in varying quantities in the glacial drift, which was derived largely from the rocks that it overlies. Recharge to bed- rock aquifers by water that moves through the surficial aquifer system also might contribute ions to the water in the bedrock aquifers.

Other minerals present in the rocks of the basin are silicates, which have minimal solubility. These minerals contribute only small quantities of dissolved ions to the ground water.

A summary of 21 representative chemical analyses of water from the limestone, Camillus, and Lockport aquifers is shown in figure 95, along with a diagram that shows the average percentage of principal constituents in the water. The water generally is very hard and of the calcium magnesium bicarbonate sulfate type. The large concentrations of sulfate indicate that the water has been in contact with gypsum. Similarly, the large concentrations of chloride indicate the presence of halite in the aquifer.

The distribution of sulfate, chloride, and hardness in water from bedrock aquifers in the Lake Erie-Niagara River Basin is shown in figures 96, 97, and 98. Concentrations of sulfate and chloride ions, as well as the hardness of the water, are greatest in the northern part of the basin in and near the area where the Camillus aquifer forms the bedrock surface, especially along Tonawanda Creek. The larger concentrations and hardness values in this area probably result partly from dissolution of gypsum and halite in the Camillus aquifer and partly because Tonawanda Creek is a principal discharge area for the regional ground-water flow system. The concentrations of all constituents in ground water tend to be larger in this discharge area, which is the end point of long flow paths.

Because much of the ground water contains sulfate and chloride in excess of 250 milligrams per liter, the quality of the water places a definite limitation on its usefulness. Similarly, excessive hardness requires that the water be treated for most uses.

In addition to the general unsuitable quality of water from shallow wells in areas where the three aquifers form the bedrock surface, the quality of water from aquifers throughout the Central Lowland and the Appalachian Plateaus Provinces generally deteriorates with depth because of limited ground-water circulation and the widespread presence of salt and gypsum beds.


The upper Housatonic River Basin is located mostly in Berkshire County in western Massachusetts (figs. 99, 85), and in a small area in Columbia County, N.Y. The basin is bounded by the drainage divide of the Housatonic River. The river drains southward through Connecticut to Long Island Sound. The basin straddles a part of a discontinuous outcrop of carbonate rocks that extends from southeastern New York and southwestern Connecticut northward along the borders of New York, Connecticut, Massachusetts, and Vermont to the Canadian border. The carbonate rocks are exposed in valleys of the Taconic and the Green Mountains and consist largely of limestone, dolomite, and marble (fig. 100). The carbonate rocks exposed in the upper Housatonic River Basin generally are geologically and hydrologically typical of the entire belt of carbonate rocks in eastern New York and western Vermont, Massachusetts, and Conncecticut, and are presented here as an example of the more extensive area.


Bedrock in the upper Housatonic River Basin consists of limestone, dolomite, and marble, as well as schist, quartzite, and gneiss (fig. 99). During various periods of geologic history, these rocks have been deformed by tilting, folding, and faulting to such a degree that their overall formation attitudes can be determined only by detailed geologic mapping. The deformation processes have caused partings along bedding planes and have created many joints, fractures, and faults, which now constitute the major water-yielding openings in the rocks.

Carbonate rocks are present in the low central part of the upper Housatonic River Valley. They consist of limestone, dolomite, and marble of Cambrian and Ordovician ages and limestone of Precambrian age (thin beds in the eastern part of the basin). As shown in figure 101, the carbonate rocks are faulted and slightly folded and are in fault contact with other rock types in several areas. The carbonate rocks are bounded on the west by quartz-mica schist with some garnetiferous schist.

The carbonate rocks are bounded on the east by quartzitic rocks that consist of quartzite, quartzite conglomerate, and feldspathic quartzite with some mica schist, and by gneissic rocks that are mostly granite-biotite gneiss with some micaceous schist and quartzite. There are numerous faults in these quartzitic and gneissic rocks.


Ground water is stored in and transmitted nearly exclusively by secondary openings in the carbonate-rock aquifer and in other rock types in the upper Housatonic River Basin. Faulting is especially important where the fault zones consist of broken rock fragments. Such zones commonly provide a large, extremely permeable conduit for water that extends to land surface and interconnects with fractures, joints, and bedding planes at depth.

In places where water circulates freely in carbonate-rock areas, such as along fault zones, dissolution may occur along openings such as joints and bedding planes, increasing and enlarging conduits for water movement. Thus, wells completed in a fault zone or solution cavity might yield large quantities of water, whereas wells a short distance away that penetrate undissolved carbonate rocks might yield virtually no water. Wells in fractured carbonate or crystalline rocks generally yield small to moderate quantities of water, depending on the number and interconnection of fractures penetrated by the well.

Ground-water movement in the basin is not well documented. However, numerous studies indicate that the size and number of openings in fractured rocks decrease with depth; this is assumed to be true in the upper Housatonic River Basin. Therefore, the ground-water flow system extends to a finite depth, possibly between 300 and 500 feet, below which fractures are sealed or sparse and flow is minimal.

Regional ground-water flow originates in the uplands (divide areas) where recharge from precipitation enters the aquifer either directly or percolates downward into the aquifer through a thin cover of glacial drift. The water then moves in response to a hydraulic gradient downward and laterally toward the Housatonic River, where it finally moves upward to be discharged into the river. Superimposed on this regional flow system are local, shallow flow systems that discharge to smaller tributary streams.

The actual flow system may differ somewhat from the theoretical flow system. Movement of ground water is subject to the random size and distribution of openings in the rocks. The water might follow irregular paths as it flows downward and laterally through the carbonate rocks. The flow system is generally unconfined; however, artesian (confined) conditions can occur. Wells that penetrate fractures filled with water under sufficient pressure to force the water upward to an altitude above that of the top of the well will flow. Similarly, a network of solution-enlarged openings in carbonate rocks might contain water under confined conditions and produce flowing wells, as well as springs where water under hydraulic pressure flows upward or laterally to the land surface. Unconsolidated glacial deposits that are thick or have minimal permeability also might confine the water in the carbonate rocks.

Yields of carbonate rocks and other rock types shown in figure 99 are comparable because each rock type represents a fractured-rock aquifer. Wells completed in carbonate rocks, with a median yield of about 9 gallons per minute, have yields that are approximately one-half the median yields of wells completed in gneissic rocks. Yields of wells completed in carbonate rocks, however, range from less than 1 to about 1,400 gallons per minute, which is a much greater range than that reported for the other rock types because yields of wells completed in carbonate-rock aquifers generally are a function of the degree of dissolution that has occurred. In some areas of extreme dissolution, yields of nearly all wells are large, but in other areas where dissolution apparently has not occurred, yields of most wells are minimal.


Water in the carbonate-rock aquifer in the upper Housatonic River Basin generally is suitable for most uses but contains large concentrations of calcium and magnesium compared to water in other rock types. This results in a moderately hard to very hard water. Bicarbonate concentrations also are large in water from the carbonate-rock aquifer. Each of these factors contributes to large concentrations of dissolved solids, which is reflected by a large specific conductance of the water. Other dissolved constituents in water from the carbonate-rock aquifer generally are in the same range as constituents in water from other rock types. This is because of the slightly mineralized precipitation that recharges all the bedrock aquifers and the similar mineralogy of the glacial deposits through which the recharge percolates before entering the aquifers.

The distribution of dissolved-solids concentrations and hardness in ground water of the upper Housatonic River Basin is shown in figure 99 and is an indication of water quality. The least mineralized and softest ground water is present around the periphery of the basin, which is in the recharge areas that generally are underlain by schist, quartzite, and gneiss that are only slightly soluble. As the water moves through these rocks toward carbonate rocks in the center of the basin and then southward, mineralization increases. The most mineralized water obtained from the carbonate-rock aquifer was from a well near the Massachusetts-Connecticut State line.

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