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feat: add epi info dataset

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@@ -25414,7 +25414,274 @@ and precipitate. Precipitation and evaporation vary with latitude and their rela
25414
  to the global wind belts. The trade winds, for example, are initially cool,
25415
  but they warm up as they blow toward the Equator. These winds pick up moisture
25416
  from the ocean, increasing ocean surface salinity and causing seawater at the surface to sink.
25417
- When the trade winds reach the Equator, they rise, and the water vapour in them condenses and forms clouds. Net precipitation is high near the Equator and also in the belts of the prevailing westerlies, where there is frequent storm activity. Evaporation exceeds precipitation in the subtropics, where the air is stable, and near the poles, where the air is both stable and has a low water vapour content because of the cold. The Greenland Ice Sheet and the Antarctic Ice Sheet formed because the very low evaporation rates at the poles resulted in precipitation exceeding evaporation in these local regions.
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
 
25418
 
25419
  The EPI - Environmental Performance Index EPI provides a data-driven
25420
  summary of the state of sustainability around the world. Using 40 performance
 
25414
  to the global wind belts. The trade winds, for example, are initially cool,
25415
  but they warm up as they blow toward the Equator. These winds pick up moisture
25416
  from the ocean, increasing ocean surface salinity and causing seawater at the surface to sink.
25417
+ When the trade winds reach the Equator, they rise, and the water vapour in them condenses and forms clouds.
25418
+ Net precipitation is high near the Equator and also in the belts of the prevailing westerlies, where there is
25419
+ frequent storm activity. Evaporation exceeds precipitation in the subtropics, where the air is stable,
25420
+ and near the poles, where the air is both stable and has a low water vapour content because of the cold.
25421
+ The Greenland Ice Sheet and the Antarctic Ice Sheet formed because the very low evaporation rates at the
25422
+ poles resulted in precipitation exceeding evaporation in these local regions.
25423
+
25424
+ Water vapour and precipitation
25425
+ As noted above, water exists in the atmosphere in gaseous form. Its liquid form,
25426
+ either as water droplets in clouds or as rain, and its solid form, as ice crystals in clouds,
25427
+ snowflakes, or hail, occur only momentarily and locally.
25428
+
25429
+ Water vapour performs two major functions: (1) it is important to the radiation balance of
25430
+ Earth, as its presence keeps the planetary surface warmer than would otherwise be the case,
25431
+ and (2) it is the principal phase of the ascending part of the water cycle.
25432
+
25433
+ The mass of water vapour in the atmosphere, which represents only 0.001 percent of the
25434
+ hydrosphere, is highest in the tropics and decreases toward the poles. At a mean temperature
25435
+ of Earth’s surface of 15 °C (59 °F), the partial pressure of water vapour at equilibrium with
25436
+ pure water is 0.017 atmosphere. The addition of salts to pure water lowers its vapour pressure.
25437
+ The equilibrium, or saturation, water vapour pressure of a saturated solution of sodium chloride
25438
+ is 22 percent lower than that of pure water. Precipitable water vapour has, on the average,
25439
+
25440
+ a vapour pressure of 0.0025 atmosphere, which amounts to 15 percent of the saturation vapour
25441
+ pressure. The ratio of observed water vapour pressure to the saturation vapour pressure at
25442
+ the same temperature is the relative humidity of the air. Thus, the mean relative humidity
25443
+ of the atmosphere is only 15 percent, a value that is low in human terms, in which
25444
+ levels of 50 to 60 percent are preferred for maximum comfort. The relative humidity
25445
+ of the air, however, varies greatly from one geographic region to another and also
25446
+
25447
+ vertically in the atmosphere. Atmospheric water vapour decreases rapidly with increasing
25448
+ altitude relative to its surface value. The amount of water required to saturate a volume
25449
+ of air depends on the temperature of the air. Air at high temperature can hold more water
25450
+ vapour at saturation than can air at low temperature. Because the temperature of the lower
25451
+ atmosphere (the troposphere) decreases rapidly with increasing altitude to about 15 km (about 9 miles),
25452
+ the upper levels of the troposphere contain little water vapour; most of the vapour is
25453
+ found within a few kilometres of Earth’s surface. The average relative humidity of tropospheric
25454
+ air is about 50 percent. Above 15 km, water vapour is essentially frozen out of the atmosphere,
25455
+ amounting to less than 0.1 percent of its concentration at Earth’s surface.
25456
+
25457
+ Aside from temperature, other factors determine the water vapour content of the air and
25458
+ are particularly important in the lower troposphere. These factors include local evaporation
25459
+ and the horizontal atmospheric transportation of moisture, which varies with altitude, latitude,
25460
+ season, and topography. During a period of 10 days (i.e., the residence time of water in the atmosphere),
25461
+ horizontal eddy turbulence may disperse vapour over distances up to 1,000 km (600 miles).
25462
+
25463
+ When a mass of air at Earth’s surface is exposed to a body of water, it gains water by evaporation or
25464
+ loses water by precipitation, depending on its relative humidity. If the air is undersaturated, with a
25465
+ relative humidity of less than 100 percent, it gains water vapour because the rate of
25466
+ evaporation exceeds the rate of condensation. If the air is supersaturated, with a relative
25467
+ humidity greater than 100 percent, the air mass loses water vapour because the rate of precipitation
25468
+ exceeds that of evaporation. This interaction between air masses and surface water bodies drives the
25469
+ atmosphere toward a state of saturation, which is not achieved for the entire atmosphere because of
25470
+ the variability in weather and because not all air masses are in contact with water bodies. In general,
25471
+ the level of atmospheric water vapour is higher in the summer, since temperatures are higher at this
25472
+ time of year. Also, atmospheric water vapour content is higher near the source of moisture than in
25473
+ distant regions. Over the oceans, the air is almost always near saturation, whereas over the deserts,
25474
+ where the supply of moisture is limited, the air is far below water vapour saturation values. In most
25475
+ cases, atmospheric water vapour content decreases inland over continents, but this decrease
25476
+ is modified by rainfall conditions, by the presence or absence of high mountains, large lakes, extensive
25477
+ forests, and swamps, and by the prevailing wind directions. Horizontal winds and air mass movements
25478
+ transfer water vapour from the ocean to the land. Although the processes are not completely separable,
25479
+ the horizontal transfer of water vapour seldom causes the vapour to undergo condensation, whereas vertical
25480
+ movements are most important in the condensation process.
25481
+
25482
+ Condensation depends strongly on the average temperature of Earth’s surface because the water
25483
+ vapour content of the air is strongly dependent on temperature. In figures that show the states of
25484
+ water as a function of the variables of pressure and temperature, the slope of the phase boundary
25485
+ between liquid water and water vapour is positive, implying that with increasing temperature the
25486
+ air at equilibrium will hold increasing amounts of water vapour. Cooling or mixing of this air
25487
+ results in condensation of the vapour and precipitation as water droplets or as ice crystals if
25488
+ the air temperature is below 0 °C (32 °F). When first formed, the water droplets or ice crystals
25489
+ are very small, on the order of 10−2 to 10−3 cm (0.004 to 0.0004 inch) in diameter, and they
25490
+ float freely in the atmosphere. In large quantities, these water droplets and ice crystals
25491
+ produce a cloud. All clouds are formed as a result of cooling below the dew point,
25492
+ the temperature at which condensation begins when air is cooled at constant pressure and constant
25493
+ water vapour content. When the droplets or crystals coalesce to a size of about 10−2 cm (0.004 inch)
25494
+ in diameter, they become heavy enough to fall as raindrops or snowflakes. Hailstones measure about
25495
+ 10−1 cm (0.04 inch) in diameter or much larger. Water vapour condensing in the atmosphere contains
25496
+ strongly soluble salts (mostly of oceanic origin), weakly soluble or insoluble solids (dust),
25497
+ and dissolved gases. The dust and sea salt aerosol particles in the air may act as sites of
25498
+ condensation by serving as nuclei for bringing initially a few water molecules together and inducing condensation from supersaturated air.
25499
+
25500
+ Distribution of precipitation
25501
+ Precipitation falling toward Earth’s surface may suffer several fates. It may be evaporated
25502
+ during its fall or after it reaches the ground surface. If the surface is covered with dense
25503
+ vegetation, much of the precipitation may be held on leaves and plant limbs and stems. This
25504
+ process is termed interception and may result in little water reaching the ground because
25505
+ the water may be directly evaporated from plant surfaces back into the atmosphere. If
25506
+ precipitation reaches the ground in the form of snow, it may remain there for some time.
25507
+ On the other hand, if precipitation falls as rain, it may evaporate, infiltrate the soil,
25508
+ be detained in small catchment areas, or become overland flow—a form of runoff. Overland
25509
+ flow (Ro) may be expressed in terms of intensity units, water depth per unit of time (e.g.,
25510
+ centimetres per hour, or inches per hour), as Chemical equation.
25511
+
25512
+ where P is precipitation rate and I is infiltration rate (rate of entry and downward movement
25513
+ of water into the soil profile). Infiltration rate will equal precipitation rate until the
25514
+ limit of the infiltration rate, or infiltration capacity, is reached. Soil infiltration rates
25515
+ are usually high at the beginning of a rain preceded by a dry spell and decrease as the rainfall
25516
+ continues. This change in rate is due to the clogging of soil pores by particles brought from
25517
+ above by the infiltrating rain and to the swelling of colloidal soil particles as they absorb
25518
+ water. Thus, rapid decreases in infiltration rates during a rain are more likely to occur in
25519
+ clay-rich soils than in sandy soils.
25520
+
25521
+ Between rainfall periods, water held in the soil as moisture is gradually lost by direct
25522
+ evaporation or by withdrawal by plants. Evaporation into the open atmosphere occurs at the
25523
+ surface of the soil, and the soil dries progressively downward with time. Water vapour in
25524
+ the soil diffuses upward, replenishing the evaporated water, and in turn is evaporated.
25525
+ The pumping of air into and out of the soil by atmospheric pressure changes enhances the
25526
+ movement of soil moisture upward. It has been shown that evaporation of a water droplet
25527
+ in the free atmosphere, and to a first approximation in various soil atmospheres, is
25528
+ proportional to the droplet surface area 4πr2 (square centimetres, where r is the radius
25529
+ of the droplet), the diffusional flux of water at the droplet surface, and the transfer
25530
+ of heat as the droplet evaporates. The equation for the rate of shrinkage of a water
25531
+ droplet due to evaporation is Chemical equation.
25532
+
25533
+ where dr/dt is the rate of change in the radius of the water droplet (centimetres per second),
25534
+ D is the diffusion coefficient of water vapour in air (cubic centimetres per second), ρvo is
25535
+ the equilibrium vapour concentration at the droplet surface, Sp is the degree of undersaturation
25536
+ of water vapour in the environment, ρL is the density of liquid water (grams per cubic centimetre),
25537
+ and X is a dimensionless parameter depending on D, ρvo, temperature, the heat of evaporation of
25538
+ water vapour, the coefficient of thermal conductivity of air, and the spherical coordinate system
25539
+ necessary to define processes occurring to a spherical water droplet. Water droplets
25540
+ shrink—dr/dt < 0, evaporate—when the water vapour concentration in the environment
25541
+ (atmosphere or soil atmosphere) is less than the saturation water vapour concentration
25542
+ at the droplet surface. They grow—dr/dt > 0, condense—when the converse is true in the
25543
+ free atmosphere. The term dr/dt has negative values for evaporation and positive ones
25544
+ for condensation. Use of this equation shows, as an example, that it would take 23 minutes
25545
+ for a water droplet to shrink (evaporate) in size from 50 to 5 micrometres (0.002 to 0.0002 inch)
25546
+ in air at 10 °C (50 °F) and a water vapour undersaturation of 1 percent.
25547
+
25548
+ Besides simple evaporation of water from soils, water is also returned to the atmosphere by
25549
+ transpiration in plants. Plants draw water from soil moisture through their vast network of
25550
+ root hairs and rootlets. This water is carried upward through the plant trunk and branches
25551
+ into the leaves, where it is discharged as water vapour. The term evapotranspiration is used
25552
+ in climatic and hydrologic studies to include the combined water loss from Earth’s surface
25553
+ resulting from evaporation and transpiration. The maximum possible evapotranspiration is
25554
+ termed potential evapotranspiration and is governed by the available heat energy. It is
25555
+ taken as equal to evaporation from a large water surface and is generally much less than
25556
+ actual evapotranspiration. Actual evapotranspiration is never greater than precipitation
25557
+ except on irrigated land because of percolation of water into groundwater bodies and surface runoff.
25558
+
25559
+ The soil moisture zone gains water by precipitation and infiltration and loses water by
25560
+ evapotranspiration, overland flow, and percolation of water downward due to gravity
25561
+ into the groundwater zone. The contact between the groundwater zone (phreatic zone) and
25562
+ the overlying unsaturated zone (vadose zone) is called the groundwater table. The water
25563
+ balance equation for change of moisture storage in a soil is given as Chemical equation.
25564
+
25565
+ where S is storage, P is precipitation, E is evaporation, and R is surface runoff plus
25566
+ percolation rate into the groundwater zone; all terms are in units of length per unit
25567
+ of time (e.g., millimetres per day, centimetres per month). In humid midlatitude climates
25568
+ where a strong contrast between winter and summer temperatures exists, there is an annual
25569
+ cycle of the water content of soils. The annual cycle of moisture in soil in Ohio, U.S.,
25570
+ for example, demonstrates the processes controlling soil moisture. Of special importance is
25571
+ the fact that the soils are saturated in this temperate climate in the spring, and the
25572
+ evaporation rate is low because of the low input of radiant energy from the Sun. By contrast,
25573
+ in the summer, evaporation increases because of increasing solar radiation, and with the growth
25574
+ of plants so does transpiration. Soil moisture is reduced to very low levels at this time of year.
25575
+
25576
+ Groundwaters and river runoff
25577
+
25578
+ The term R in the water balance equation for change of soil moisture storage above represents
25579
+ groundwater and river runoff losses from the soil moisture zone. Water percolates from the soil
25580
+ moisture zone through the unsaturated (vadose) zone to the water table. Flow through the
25581
+ unsaturated zone is complicated. After a rainfall, water may form a nearly continuous phase
25582
+ in pores in this zone, but, with drying, the last amount of water is held in clusters at points
25583
+ of contact of solid grains and as thin films on solid surfaces. The flow paths of water become
25584
+ more tortuous, and the water-conducting properties decrease rapidly. Structured soils and fractured
25585
+ rock in the vadose zone may act as conduits for fluids to reach the water table. Because of the complex
25586
+ geometry of water contained in the unsaturated zone, the properties of water are expressed by means of
25587
+ empirical relationships. Darcy’s law, derived in 1856 from experimentation by the French engineer Henri
25588
+ Darcy, permits quantification of water flow through porous media. The law states that the rate of
25589
+ flow Q of a fluid through a porous layer of medium (e.g., a sand bed) is directly proportional to the
25590
+ area A of the layer and to the difference Δh between the fluid heads at the inlet and outlet faces of
25591
+ the layer and is inversely proportional to the thickness L of the layer. Expressed analytically, Chemical equation.
25592
+
25593
+ where K is a constant characteristic of the medium.
25594
+ The term K for a porous rock medium is the volume of fluid of unit viscosity passing
25595
+ through a unit cross section of the rock in unit time under the action of a unit pressure
25596
+ gradient; this characteristic is called permeability. The permeability of a rock is
25597
+ dependent on the geometric properties of the rock, such as porosity, shape and size
25598
+ distribution of constituent rock grains, and degree of cementation of the rock.
25599
+ Permeabilities of rocks vary greatly. Unconsolidated sands may have permeabilities
25600
+ measured in hundreds of darcys, whereas consolidated sands that will transmit reasonable
25601
+ amounts of fluid have permeabilities of 0.01 to 1 darcy. A rough idea of the meaning of
25602
+ one darcy of permeability (which equals 9.869 × 10−12 square metre [1.261 × 10−11 square foot])
25603
+ can be obtained by imagining a cube of sand one foot on a side. If the sand has a permeability
25604
+ of one darcy, approximately one barrel of water per day will pass through the one-foot cube with
25605
+ a one-pound pressure head. The general equation of Darcy can be modified to express flow in both
25606
+ the unsaturated zone and the saturated groundwater zone.
25607
+
25608
+ Groundwater is constantly in motion. When a lake or stream intersects the groundwater table,
25609
+ groundwater communicates directly with these bodies of water. If the groundwater table is higher
25610
+ than the stream or lake level, a pressure head will develop such that the groundwater flows
25611
+ into the water body; conversely, if the groundwater table is lower than the river or lake level,
25612
+ the pressure gradient induces flow into the groundwater. Most groundwater ultimately reaches the
25613
+ channels of surface streams and rivers and flows to the sea. On the average, groundwater contributes
25614
+ to total river runoff about 30 percent of its water on a global basis.
25615
+
25616
+ Water runoff from the land surface is that part of precipitation which eventually appears in
25617
+ perennial or intermittent surface streams. Streamflow-generation mechanisms have been studied
25618
+ for several decades, and there is now considerable knowledge regarding rainfall runoff processes
25619
+ and their controls. This understanding is the result of both careful observations from field
25620
+ experiments and the heuristic simulations of hypothetical realities with rigorous mathematical models.
25621
+ The discharge measured at the downstream end of a channel reach is supplied by channel inflow at the
25622
+ upstream end of the reach and by the lateral inflows that enter the channel from the hillslope along
25623
+ the reach. The lateral inflows may arrive at the stream in one of three forms: (1) groundwater flow,
25624
+ (2) subsurface storm flow, or (3) overland flow.
25625
+
25626
+ Groundwater flow provides the base flow component of streams that sustains their flow between
25627
+ storms. The “flashy” response of streamflow to individual precipitation events may be ascribed
25628
+ to either subsurface storm flow or overland flow. Under intense rainfall events during which
25629
+ the surface soil layer becomes saturated to some depth, water is able to migrate through “preferred pathways”
25630
+ rapidly enough to deliver contributions to the stream during the peak runoff period. The conditions for
25631
+ subsurface storm flow are quite restrictive. The mechanism is most likely to be operative on steep, humid,
25632
+ forested hillslopes with very permeable surface soils.
25633
+
25634
+ Overland flow is generated at a point on a hillslope only after surface ponding takes place.
25635
+ Ponding cannot occur until the surface soil layers become saturated. It is now widely recognized
25636
+ that surface saturation can occur because of two quite distinct mechanisms—specifically, Horton overland
25637
+ flow (named for American hydraulic engineer and hydrologist Robert E. Horton) and Dunne overland flow
25638
+ (named for British hydrologist Thomas Dunne).
25639
+
25640
+ The former classic mechanism is for a precipitation rate that exceeds the saturated hydraulic
25641
+ conductivity of the surface soil. A moisture content versus depth profile during such a rainfall
25642
+ event will show moisture contents that increase at the surface as a function of time. At some point
25643
+ in time the surface becomes saturated, and an inverted zone of saturation begins to propagate downward
25644
+ into the soil. It is at this time that the infiltration rate drops below the rainfall rate and overland
25645
+ flow is generated. The time is called the ponding time. The necessary conditions for the generation of
25646
+ overland flow by the Horton mechanism are (1) a rainfall rate greater than the saturated hydraulic
25647
+ conductivity of the soil and (2) a rainfall duration longer than the required ponding time for a given
25648
+ initial moisture profile. Horton overland flow is generated from partial areas of the hillslope where
25649
+ surface hydraulic conductivities are lowest.
25650
+
25651
+ In Dunne overland flow, the precipitation rate is less than the saturated hydraulic conductivity,
25652
+ and the initial water table is shallow or there is a shallow impeding layer. Surface saturation occurs
25653
+ because of a rising water table; ponding and overland flow occur at a time when no further soil moisture
25654
+ storage is available. The Dunne mechanism is more common to near-channel areas. Dunne overland flow is
25655
+ generated from partial areas of the hillslope where water tables are shallowest. Both Horton and Dunne
25656
+ mechanisms result in variable source areas that expand and contract through wet and dry periods.
25657
+
25658
+ Total river discharge and the chemistry of the discharge vary from continent to continent; some
25659
+ continents are wetter and some drier than the world average, but the deviations are not extreme.
25660
+ The runoff per unit area from Asia and Europe is almost exactly equal to the world average; it is a
25661
+ little lower in Africa and North America; and it is considerably higher in South America. Antarctica
25662
+ is frozen and Australia is arid, and so they contribute little runoff. Also, since their areas are
25663
+ relatively small, they do not affect the global runoff average significantly. The waters draining
25664
+ the continents have quite different chemistries; those from Europe are very rich in calcium and
25665
+ bicarbonates, whereas those from Africa and South America are not. North American and Asian rivers
25666
+ re somewhat intermediate in their concentrations of these dissolved constituents. Such differences
25667
+ in composition reflect a variety of factors, including runoff, temperature, and relief, but certainly
25668
+ the bulk composition of the continental rocks in contact with these waters and their underground sources
25669
+ play a major role. The surface rocks of Europe are rich in carbonates, and those of South America are not;
25670
+ the latter are dominated by sediments rich in silicate minerals.
25671
+
25672
+ The chemistry of groundwater and river runoff is being modified by human activities on a global
25673
+ scale. The natural dissolved riverine input of major constituents to the oceans already has been
25674
+ increased by more than 10 percent because of human activities. In the case of sodium, chlorine,
25675
+ and sulfate, the increases are as high as 30 percent. In the United States alone, total water use
25676
+ is equivalent to one-third of total runoff, with about 2 percent of the water used coming from
25677
+ underground wells. In the southwestern region of the country, water supplies have been tapped heavily
25678
+ and in some areas have been exhausted with no hope of replacement. This extensive use of fresh waters in
25679
+ the United States and throughout the globe makes them particularly susceptible to pollution. Leachates
25680
+ from fertilizers, herbicides, and pesticides are found in some freshwater bodies; toxins or excessive
25681
+ amounts of certain inorganic or organic chemicals are present; radioactive elements have been detected;
25682
+ and some surface water bodies have had their salinities increased dramatically, rendering them useless
25683
+ for human consumption. It is therefore imperative that countries closely monitor the use of freshwater
25684
+ systems and promote their conservation.
25685
 
25686
  The EPI - Environmental Performance Index EPI provides a data-driven
25687
  summary of the state of sustainability around the world. Using 40 performance